Initiation and Evolution of the Aleutian Arc and Comparison to the Izu-Bonin-Mariana Arc

INITIATION AND EVOLUTION OF THE ALEUTIAN ARC AND COMPARISON TO THE IZU-BONIN-MARIANA ARC

A Dissertation Presented to the Faculty of the Graduate School of Cornell University in Partial Fulfillment of the Requirements for the Degree of Doctor of Philosophy

by Ashley Kaye Tibbetts December 2018

INITIATION AND EVOLUTION OF THE ALEUTIAN ARC AND COMPARISON TO THE IZU-BONIN-MARIANA ARC Ashley Tibbetts, Ph.D. Cornell University 2018

Subduction zones play an important part in shaping our planet. Yet, we cannot presently observe the mechanisms by which they initiate or their subsequent evolution. It is for this reason that we have chosen the Aleutian and Izu-Bonin-Mariana (IBM) Arcs to investigate the initiation of subduction around the Pacific Rim at ~50 Ma. These arcs initiated at approximately the same time, have older rocks exposed for study, and exhibit a variety of subduction characteristics along their lengths, enabling us to investigate not only their initiation, but their evolution through time. The Aleutian Arc in particular changes convergence angle and chemistry rather significantly along arc, allowing us to investigate the effect of the oblique subduction in the western Aleutians on the geochemistry. New geochemical and geochronologic data on Attu and

Kiska have enabled us to gain a more complete understanding of the processes occurring in this part of the arc. Through these investigations, the following conclusions have been reached:

1) Initial volcanism on Attu was tholeiitic; it became more incompatible element-depleted in

the north due to incipient back-arc rifting, with a more extensional signature in the north and

a stronger subduction signature in the south; rifting magmatism was characterized by

fractional crystallization and assimilation at 16 Ma; and Attu abruptly switched to calc-

alkaline subduction volcanism at ~6 Ma.

2) Kiska’s geochemistry is intermediate between that of Attu (western Aleutians) and Adak

(central Aleutians), which coincides with its placement on the ridge. Volcanism shifted

northward and transitioned to smaller volume volcanism and a more compressive

environment, eventually leading to the present calc-alkaline regime.

3) Based on their geochemistry Rat Island is part of the Aleutian Arc and is not

geochemically related to Bowers Ridge. Rat Island transitioned from tholeiitic to calc- alkaline at ~15 Ma due to deeper melting, a more water-rich source, and a shift from a transtensional to a transpressional tectonic environment due to block rotation and changes in plate motion. Rat Island experienced the same northward shift in volcanism as the rest of the arc, and its geochemistry reflects its placement on the Aleutian Ridge.

4) By comparing the IBM and Aleutian arcs, we investigated variations in the types of subduction initiation occurring around the Pacific at ~50 Ma and how this affected the evolution of these arcs. The IBM Arc experienced a decreasing flux of sediments with time.

The Bonin had the largest contribution of sediment, followed by the older Mariana samples, the modern Mariana samples, and the Izu and Aleutian samples. The older IBM samples experienced more variable degrees of partial melting. The Mariana fore-arc basalts and the mid-ocean ridge (MORB-like) Attu samples represent the depleted end-members of the two arcs. The Mariana segment experienced the largest degree of partial melting, followed by the

Bonin segment, and the Izu segment. The fluid flux remained relatively stable with time. The fluid flux is greatest in the Aleutians, followed by the Mariana segment, the Bonin segment, and the Izu segment. Most of the fluid was derived from oceanic crust rather than sediment.

BIOGRAPHICAL SKETCH

Ashley was born in Sanbornton New Hampshire and graduated from Winnisquam

Regional High School. She earned her bachelors of arts degree in earth science and astronomy from Boston University in 2007. She graduated from the University of Nevada, Las Vegas in

2010 with a Masters in geoscience. Her Master’s thesis was entitled “Petrogenesis of the

Greenwater Range: Comparison to the Crater Flat Volcanic Field and Implications for Hazard

Assessment.” She is currently teaching high school chemistry at Catholic Memorial School in

Massachusetts and scientific editing with Accdon LLC.

ACKNOWLEDGEMENTS

I would like to thank Bill White for his infinite patience and support through the long writing process and to Suzanne Kay for her involvement and dedication throughout the project and her dedication to reading each draft. Thank you to Robert Kay and Steve Squyres for their comments and to Brian Jicha and Allen Schaen at the University of Wisconsin for sharing the new geochronology data used here. Thank you also to Brian Jicha, Suzanne Kay and Allen

Schaen for participating in new field studies on Kiska, Rat (Haladax), and Adak Islands during the tenure of this study and to Robert Kay, Suzanne Kay, James Rubenstone, Gene Yogodzinski,

Gary Citron, and Steve Squyres among others for previous field work and sampling and analyses that contributed to this study. Thank you to Karen Harpp at Colgate University for her help with

the ICP analyses conducted at Colgate. Thank you to Gregg McElwee for his help with the ICP

analyses conducted at Cornell. Support for this project was provided by National Science

Foundation Earth Science grant # EAR‐1144494 to Suzanne Kay and Brad Singer and a long

series of prior NSF grants that supported prior analyses and sample collection. Additional

financial support for this project was provided by teaching assistantships from the Department of

Earth and Atmospheric Sciences at Cornell University. Thank you to Savanah Sawyer for her

help with all the forms and paperwork.

TABLE OF CONTENTS

Biographical Sketch ...... iii

Acknowledgments...... iv

Table of Contents ...... ………..vi

List of Figures ...... viii

List of Tables ...... x

Chapter 1: The Western Aleutian Islands: The History of Attu and Kiska Islands and Implications for the Initiation and Evolution of the Aleutian Arc ...... 1 INTRODUCTION ...... 1 BACKGROUND AND PREVIOUS WORK ...... 5 Tectonic History of the Bering Sea and North Pacific Ocean ...... 5 Background Geology of Attu Island ...... 8 Background Geology of Kiska Island ...... 22 NEW STUDIES OF ATTU AND KISKA ISLANDS ...... 29 Sample Collection and Analytical Methods ...... 29 Attu Island Samples ...... 42 Kiska Island Samples ...... 46 RESULTS ...... 57 Major Elements ...... 57 Trace Elements...... 69 Trace Element Ratios of Attu and Kiska Islands Compared to North Pacific MORB and Adak Island in the Central Aleutians ...... 73 Isotope Ratios...... 80 DISCUSSION ...... 85 Magmatic and Tectonic Evolution of Attu and Kiska Islands ...... 85 Attu Island ...... 85 Kiska Island ...... 96 Tholeiitic to Calc-Alkaline Evolution ...... 99 Tectonic Model for Attu and Kiska Islands ...... 111 Attu Island ...... 111 Kiska Island ...... 118 SUMMARY ...... 121 REFERENCES ...... 125

Chapter 2: Rat (Hawadax) Island in the Context of the Aleutian Arc and Bowers Ridge ...... 136 INTRODUCTION ...... 136 GEOLOGIC BACKGROUND ...... 14 Rat Island ...... 143 Bowers Ridge ...... 146 Bowers Basin ...... 148

SAMPLEING AND ANALYTICAL METHODS ...... 154 RESULTS ...... 154 Major Elements ...... 154 Trace Elements...... 161 Isotope Ratios...... 169 DISCUSSION ...... 172 Rat Island: Part of Bowers Ridge or the Aleutian Arc? ...... 172 Magmatic Evolution on Rat Island ...... 175 Evolution of the Rat Island Block Through Time ...... 181 CONCLUSIONS...... 188 REFERENCES ...... 193

Chapter 3: A Comparison of Sediment Subduction in the Aleutian and Izu-Bonin-Mariana Arcs Through Time ...... 201 INTRODUCTION ...... 201 GEOCHEMISTRY OF THE ALEUTIAN AND IZU-BONIN-MARIANA ARCS ...... 208 COMPARISON OF THE ALEUTIAN AND IZU-BONIN-MARIANA ARCS ...... 226 IMPLICATIONS FOR SUBDUCTION INITIATION AROUND THE PACIFIC RIM…...... 250 REFERENCES ...... 253

APPENDIX A: Analytical Methods ...... 263 Major Elements ...... 263 Isotope Ratios...... 265 Trace Elements...... 266 Ar/Ar Geochronology ...... 270 APPENDIX B: Attu Island Sample Descriptions ...... 271 APPENDIX C: Rat Island and Kiska Island Thin Section Descriptions ...... 280 APPENDIX D: Additional Data ...... 283

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LIST OF FIGURES Figure 1.1 Map of the Bering Sea region showing relevant structural features and islands...... 2 Figure 1.2 Map of Attu Island, legend, and sample number key ...... 10 Figure 1.3 Map of Kiska Island with legend...... 24 Figure 1.4 Chondrite normalized spider diagrams for Attu Island ...... 30 Figure 1.5 Primitive mantle normalized spider diagrams for Attu Island ...... 35 Figure 1.6 SiO2 vs. Feo/Mgo (Tholeiitic vs. Calc-alkaline) discrimination diagrams for Attu and Kiska Islands ...... 40 Figure 1. 7 SiO2 vs. K2O discrimination diagrams for Attu and Kiska Islands ...... 41 Figure 1.8Chondrite normalized spider diagrams for Kiska Island ...... 48 Figure 1.9 Primitive mantle normalized spider diagrams for Kiska Island ...... 51 Figure 1.10 SiO2 vs. La/Sm (normalized to C1 chondrite) for Attu and Kiska Islands ...... 72 Figure 1.11 Sm/Yb vs. La/Sm normalized to C1 chondrite for Attu and Kiska Islands ...... 76 Figure 1.12 La/Ta vs. Ba/La for Attu and Kiska Islands ...... 77 Figure 1.13 Th/La vs. Ba/La for Attu and Kiska Islands ...... 78 87 86 Figure 1.14 Sr/ Sr vs. εNd for Attu and Kiska Islands ...... 79 Figure 1.15 εNd vs. La/Sm (normalized to C1 chondrite) for Attu and Kiska Islands ...... 82 Figure 1.16 εNd vs. Th/La for Attu and Kiska Islands ...... 83 207 204 Figure 1.17 SiO2 vs. Pb/ Pb ...... 94 Figure 1.18 Age vs. Th/La, εNd, La/Sm (Normalized to C1 chondrite), and Ta/La ...... 98 Figure 1.19 Tectonic evolution diagram for the Bering Sea region...... 106 Figure 1.20 Diagram of the evolution of Attu...... 110

Figure 2.1 Map of the Bering Sea region showing relevant structural features and islands...... 137 Figure 2.2 Map of Rat Island ...... 142 Figure 2.3 Pictures of Rat Island...... 150 Figure 2.4 SiO2 vs. K2O discrimination diagrams ...... 156 Figure 2.5 SiO2 vs. FeO/MgO (calc-alkaline vs. tholeiitic) discrimination diagrams ...... 158 Figure 2.6 Chondrite normalized spider diagrams ...... 164 Figure 2.7 Primitive mantle normalized spider diagrams ...... 165 Figure 2.8 Sm/Yb vs. La/Sm normalized to C1 chondrite ...... 166 Figure 2.9 La/Ta vs. Ba/La ...... 167 Figure 2.10 Th/La vs. Ba/La ...... 168 87 86 Figure 2.11 Sr/ Sr vs. εNd ...... 170 Figure 2.12 εNd vs. Th/La ...... 171 Figure 2.13 SiO2 vs. Sm, La, Yb, and (La/Sm)N for the Gunners Cove and Rat Formations .....177 Figure 2.14 La vs. La/Sm for the Gunners Cove and Rat Formations ...... 178 87 86 Figure 2.15 Distance from trench vs. (Sm/Yb)N, (La/Sm)N, Ba/La, Ta/La, Th/La, Sr/ Sr, and εNd ...... 184

Figure 3.1 Map of Northwest Pacific Ocean showing relevant features ...... 203 Figure 3.2 SiO2 vs. K2O discrimination diagram ...... 211 Figure 3.3 Th/La vs. K2O ...... 213 Figure 3.4 Chondrite normalized spider diagram ...... 214 Figure 3.5 Sm/Yb vs. La/Sm normalized to C1 chondrite ...... 217 Figure 3.6 Th/La vs. Ba/La ...... 218

viii

87 86 Figure 3.7 Sr/ Sr vs. εNd mixing diagram ...... 220 Figure 3.8 εNd vs. Ba/La ...... 223 Figure 3.9 La/Sm (normalized to C1 chondrite) vs. εNd ...... 224 Figure 3.10 Th/La vs. εNd ...... 225 Figure 3.11 Primitive mantle normalized spider diagram of sediments ...... 228 Figure 3.12 Age vs. Th/La, εNd, La/Sm (Normalized to C1 chondrite), and Ta/La ...... 249

LIST OF TABLES

Table 1.1: 40Ar/39Ar Ages for Attu and Kiska Islands ...... 17 Table 1.2: Kiska Island Sample Locations and descriptions ...... 54 Table 1.3: Major Elements for Attu Island Samples ...... 60 Table 1.4: Major Elements for Kiska Island Samples ...... 61 Table 1.5: Trace Elements for Attu Island Samples ...... 63 Table 1.6: Trace Elements for Kiska Island Samples ...... 66 Table 1.7: Sr and Nd Isotopes for Kiska and Attu Islands...... 84

Table 2.1: Rat Island 40Ar/39Ar Ages ...... 145 Table 2.2: Rat Island Sample Descriptions ...... 149 Table 2.3: Rat Island Major Elements ...... 155 Table 2.4: Rat Island Trace Element Concentrations (ppm)...... 159 Table 2.5: Rat Island Isotopes ...... 169

Table 3.1: Correlation coefficients for various trace element ratios versus εNd ...... 234

Chapter 1: The Western Aleutian Islands: The History of Attu and Kiska Islands and Implications for the Initiation and Evolution of the Aleutian Arc

INTRODUCTION

Subduction of oceanic crust is an integral and important part of plate tectonics, which is

the process by which the Earth loses its internal heat and which continually reshapes the Earth’s

surface. Subduction zones are also important areas of crust-to-mantle and mantle-to-crust mass

transfer. Subduction is often accompanied by major earthquakes and explosive volcanic

eruptions, and it is the aspect of plate tectonics that most directly affects human society.

Subduction zones now surround much of the Pacific Ocean, resulting in the so-called Ring of

Fire, but was it always so? While subduction and its connection to earthquakes is relatively well

understood, how new subduction zones initiate and how they evolve remains poorly understood.

Thus, understanding subduction zone initiation is the key to understanding the tectonic evolution of the Earth.

Though subduction has been occurring for billions of years, if not for all of Earth’s history, all presently observable island arcs are relatively young. Indeed, Gurnis et al. (2004) note that nearly half of all active subduction zones initiated in the Cenozoic. Clearly then, subduction initiation is a common and important process, but understanding subduction zone initiation is inhibited by an absence of currently initiating subduction zones. The alternative to direct observation is to study the oldest rocks in mature subduction systems. This is complicated by the fact that the older rocks erupted are subsequently buried by younger eruptions.

The Bering Sea region (Figure 1.1) has a complex tectonic history, and the margins surrounding it have changed significantly over the last ~50 Ma. A large part of the evolution of

Figure 1.1: Bering Sea Region showing various structural features and several Islands on the Aleutian Ridge.

the Aleutian Arc. Though many studies of the various parts of the Aleutian Arc have been conducted over the years, little work has been done on Kiska Island, and more information is needed to gain a more complete history of Attu Island. This study focuses on the Western

Aleutian Arc and how it has changed through time and with distance along the arc. In particular, we seek to extend our understanding of the histories of Attu and Kiska in an effort to better understand the evolution of the Western Aleutians.

The Aleutian Arc is a chain of volcanoes that have formed above the convergent margin where the Pacific Plate is being subducted beneath the North American Plate. This subduction zone is ideal for investigating questions about the initiation and evolution of arc volcanics for several reasons. First, the volcanic line has migrated northward (e.g., Kay et al., 1990) placing younger eruptive material further north rather than completely burying the older rocks. Second, glacial erosion has locally removed more recent eruptive material, exposing older Paleogene volcanic and intrusive rocks. At the same time, however, strike slip faulting and forearc subduction erosion may also have removed part of the arc, so the record of the earliest arc is likely incomplete. Third, magmatism has decreased in intensity in the Western Aleutians over the Cenozoic as the motion between the North American and Pacific plates has become more oblique, resulting in low absolute convergence rates at present (Delong et al., 1978; Krutikov et al., 2008). Along its ~3,800 km length the Alaska-Aleutian Arc goes from almost orthogonal in the east to almost completely strike-slip in the west.

Other factors also vary along the length of the arc. The Pacific lithosphere that is currently being subducted is Late Paleocene to Early Cretaceous in age (Miller et al., 2006;

Norton, 2007; Scholl, 2007) and current slab dips vary from 45° in the east to 60° in the central

Aleutians to 50° in the western Aleutians (e.g., Ruppert et al., 2007). The Aleutian ridge also

narrows from ~250 km in the east to 80 km near the Komandorsky Islands (see review in Scholl,

2007). The ~35-37 km crustal thickness throughout most of the arc (~27 km on Attu;

Janiszewski et al., 2013) is unusual for an intra-oceanic arc, and the mid- to lower-crustal seismic

velocities indicate a mafic composition (Holbrook et al., 1999; Lizarralde et al., 2002;

Shillington et al., 2004; Janiszewski et al., 2013). These changes in geophysical setting combined

with along-arc geochemistry of exposed rocks from different stages of arc history allow for investigation of the initiation and evolution of the Aleutian Arc.

The western part of the Aleutian Arc has a particularly interesting history. As a well- studied example of the western arc, the Pre-Miocene rocks on Attu (Figure 1.1) have a chemical composition that distinguishes them from those in the eastern and central Aleutian Islands (e.g.,

Rubenstone, 1984; Yogodzinski et al., 1993; Kay and Kay, 1994). Some of the older rocks on

Attu, including pillow lavas, volcanic breccias, dikes, and gabbros, have some chemical similarities to mid-ocean ridge basalt (MORB) in terms of depleted isotopic signatures and light rare earth element (LREE) depleted REE patterns. This, as well as the submarine style of eruption, hydrothermal alteration, and weak arc-related chemical signatures, contrasts with the distinctly arc-like compositions of the older rocks in the central, e.g., Adak, and eastern

Aleutians and suggests a distinct tectonic setting and magmatic evolution for the western part of the arc at that time. Nevertheless, the arc-related trace element signatures of the tholeiitic basalts within the Attu Basement Series suggest the series was formed in the vicinity of a subduction zone. Yogodzinski et al. (1993) suggested that the Attu basement series formed in a rift-related transtensional environment where “the lithosphere of the overriding plate is stretched in response to tectonic adjustments resulting from a switch from convergence to strike-slip-dominated tectonics.”

This study builds on previous studies of Attu (e.g., Gates et al., 1971; Rubenstone, 1984;

Shelton, 1986; Yogodzinski et al., 1993) with the goal of achieving a better understanding of the subduction initiation and magmatic and tectonic evolution of the Western Aleutians. New

40Ar/39Ar ages for 11 samples (Jicha, personal communication) and chemical data for 21 samples from Attu augment the geochemical and geochronologic data of previous studies.

New geochronologic (Jicha, personal communication) and geochemical data are also reported for the magmatic rocks of Kiska Island (Figure 1.1). Kiska’s location as one of the islands closest to the trench as well as the fact that it has some of the oldest reported ages

(55.3±6.7 Ma Lief Cove: Marvin and Cole, 1978; 46 Ma submarine Murray Canyon: Jicha et al.,

2006), made it an obvious candidate for potential exposures of early arc rocks. Thus, new sampling of Kiska was undertaken with the objective of sampling some of these older rocks.

Furthermore, it is one of the few westernmost Aleutian Islands with historic on-land volcanic activity. Kiska’s long history also means that the chemistry and ages of these samples are useful in understanding the evolution of the western Aleutians over an extended period. In addition, comparison of Kiska and Attu reveals how magmatism has varied, both at any given point in history and through time along the arc. It should be noted that there is subaerial volcanism on

Buldir (Scholl et al., 1976) and recent submarine volcanism west of Kiska (Yogodzinski et al.,

2015; Yogodzinski et al., 2017).

BACKGROUND AND PREVIOUS WORK

Tectonic History of the Bering Sea and North Pacific Ocean

Scholl (2007; Figures 2, 10A, 10B) described the changes in tectonic configuration that led to the initiation of the Aleutian subduction zone and the present-day configuration of the

Bering/North Pacific region. This history begins with the accretion of the Olyutorsky Arc onto the Kamchatka-Koryak Margin. This arc accretion led to the obstruction of the subduction zone and buildup of stress as the Pacific Plate continued to move north without subduction

accommodating its motion. Scholl (2007) places this event at ~45-55 Ma since it takes time for

the arc to accrete and completely stop subduction. In his model, the buildup of compressive

stress on a NW-SE axis combined with the focusing of buckling forces where the SW-striking

Alaska subduction zone met the NW striking Beringian Margin then caused the Aleutian

subduction zone to form as a seaward extension of the Alaska subduction zone along a west

trending fracture zone that extended off the Alaska Peninsula. The Aleutian subduction zone

subsequently accommodated the Pacific Plate motion, and the major stress in the Bering Sea area

was transferred to the piece of oceanic crust, Aleutia, trapped north of the Aleutian subduction

zone. This stress was caused by an event termed the North Pacific Rim orogenic stream

(Redfield et al., 2007), which is the extrusion of Alaskan lithosphere SW across the Beringian

Margin. In Scholl’s model, this SW-NE stress was accommodated by the formation of the

Shirshov and Bowers subduction zones contemporaneous with or slightly after the initiation of

the Aleutian subduction zone. Present-day Bowers Ridge has an inferred fossil trench on its

northern flank (Ludwig et al., 1971; Kienle, 1971; Rabinowitz, 1974; Cooper et al., 1981; Ben-

Avraham and Cooper, 1981; Sato et al., 2016). Though there is no evidence of a trench bordering either side of Shirshov Ridge (Ludwig et al., 1971; Kienle, 1971), one of the possible scenarios

for the tectonic history of the Bering Sea region proposed by Ben-Avraham and Cooper (1981)

includes the initial presence of a trench on Shirshov Ridge’s western side that was later

destroyed by rifting during the extension responsible for the Komandorsky Basin. However,

based on the MORB-like chemistry of dredged gabbros and zircon dating, Sukhov et al. (2011)

concluded that Shirshov Ridge is compressionally deformed Mid- to Late-Cretaceous oceanic

crust. A K-Ar age of 47±5 Ma for an amphibolite suggests that this deformation occurred around

47 Ma. Thus, although the Shirshov Ridge did not develop into a subduction zone, its structure is consistent with the direction of the Eocene compression envisioned by Scholl (2007). In addition to these ridges, spreading in the Shirshov backarc could have produced the Vitus Ridge and

Bowers Basin shortly after the capture of Aleutia by initiation of the Aleutian subduction zone.

Scholl places the initiation of the Aleutian subduction zone in or earlier than the Middle

Eocene based on the 46 Ma age (40Ar/39Ar, groundmass; Jicha et al., 2006) of a basaltic andesite

from Murray Canyon ,west of Kiska, and the formation of the Shirshov and Bowers subduction

zones at the same time or slightly after the formation of the Aleutian subduction zone. Based on

the ages in Wanke et al. (2012) for the crest of Bowers Ridge (~32 Ma), Scholl’s (2007)

assumption that Bowers Ridge and the Aleutian Ridge initiated at approximately the same time is

plausible. A new 47.1 ±3.6 Ma 40Ar/39Ar age on plagioclase from a granulite xenolith from

Kanaga (Kay et al., 2014) further confirms initiation of Aleutian volcanism by the Middle

Eocene.

Although these are the most reliable older ages, two older, less precise ages exist. One is

the K/Ar age from the southwestern coast of Kiska 55.3±6.7 Ma, which is a minimum of 48.6

Ma given the error (Marvin and Cole, 1978); and the other is a plagioclase 40Ar/39Ar age from

Adak that plateaued at 34.7 Ma, but had a higher temperature tail at 50 Ma (Rubenstone, 1984).

Rubenstone (1984) interpreted this age as a 34 Ma metamorphic event that partially reset the >50

Ma crystallization age. This would suggest that subduction had initiated and the Aleutian Ridge

had formed by 50 Ma. In addition, a 41.5+/- 2 Ma amphibole age reported by Vallier et al.

(1994) and a 40.3±2 0.1 Ma Attu basalt intruded by 30-35 Ma plutons reported by Kay et al.

(2014) indicate that the western Aleutian Ridge had built up significantly by this time. Jicha et al.

(2006) estimated a ridge building minimum magmatic addition rate of 89 km3/km/Ma, and 182

km3/km/Ma if mid-crustal plutons and cumulates and erosion are taken into account, which is several times higher than earlier estimates.

Further, Wanke et al. (2012) provide a new estimate for the timing of volcanism on

Bowers Ridge based on a 32.3±2.0 Ma age for an andesite on Bowers Ridge and a seamount age indicating that the Bowers Ridge structure was active until at least 22.2 Ma (Wanke et al., 2012).

Other constraints include the formation of the Komandorsky Basin at ~20-9 Ma (Valyasho et al.,

1993; Scholl, 2007).

Background Geology of Attu Island

As shown in Figure 1.1, Attu is the westernmost U.S. Aleutian Island and the largest of the Near Island group (Attu, Agattu, and Shemya). Kiska Island is approximately 300 km to the

southeast. Both islands are west of Bowers Ridge, southeast of Shirshov Ridge, and south of

Bowers Basin. Due to the change in subduction angle along arc, both Attu and Kiska are currently experiencing more oblique subduction than the central and eastern parts of the arc

(Figure 1.1). Although the velocity of the Pacific Plate relative to the arc increases slightly from east to west (Figure 1.1), because of the changing convergence angle, the orthogonal convergence rate decreases dramatically from the central Aleutians (~6.5 cm/yr) to the western

Aleutians (~1.2 cm/yr) (e.g., Syracuse and Abers, 2006). Slab dips below 100 km shift from 45° in the eastern Aleutians to 60° in the central Aleutians to 50° in the western Aleutians (Ruppert et al., 2007). This results in the slab being initially subducted to the east of the islands and then traveling laterally west before reaching a significant depth below the islands. The increased

transport time from additional lateral movement results in a greater slab temperature at a given depth than would be present below the central and eastern arc (Yogodzinski et al., 1993).

Capps (1934) describes Attu as being 915-1065 m in elevation with few lowlands and having cliff-like shores with little or no beach. He describes the overall nature of the island as heavily glaciated with no active volcanic vents. Gates et al. (1971) describe these cliff-like shores as being marine terraces a hundred meters high and up to 533 m on the north shore with

1-3 wave cut benches.

Gates et al. (1971) described and divided the magmatic rocks of Attu (Figure 1.2) into an older Attu Basement Series, several younger formations, and the youngest Attu Calc-Alkaline

Series. The Basement Series consists of pillow lavas and submarine volcanoclastics interbedded with a variety or marine sedimentary rocks, including argillites, sandstones, graywackes, and conglomerates. Small amounts of amphibolite (caused by an early metamorphic event that altered many of the earliest volcanic rocks), with a K-Ar age of 41.5±2 Ma (Vallier et al., 1983), have been observed on southern Attu, east of Temnac Bay in the area between Kaufman Creek and the Temnac River. Gates et al. (1971) place these amphibolites in the Basement Series. The thickness of the basement rocks is at least 2130 m and could be as much as 4570 m if erosion is taken into account. The basement rocks dip < 30° N. The fine-grained sedimentary rocks include cherts, siliceous siltstone, argillite, limestone, and tuffaceous greywacke. These rocks are characterized by thin beds, from several centimeters to several meters thick, some unconformities, deformation, and turbidity current textures. They are mostly composed of weathered pyroclastics, lavas, and tuffs. The coarser sedimentary rocks include greywackes and conglomerates and are present in profusion on the northwestern third of the island. The conglomerates are massive with little structure and are composed of detrital rock fragments,

Figure 1.2: A) Geologic map of Attu (after Gates et al., 1971, plate 80) with sample locations indicated; B) Map key; C) Table listing sample names and ages with corresponding map numbers.

172°30’E 173°15’E 53°00’N 1 2 Steller Austin 53°00’N Cove 7 Cove 6 8 18 3 4 5 19

10 12 16 17 20 9 23 Chichagof 11 24 21 31 22 32 Harbor 13 15 25 30 26 29 33 14 34 Etienne 27 Bay 28 35 39 36 37 Abraham Sarana Bay Bay 38 41 40 Attu 49 42 43 48 50 51 53 54 52 57 58 Massacre 59 64 47 Island 61 Bay 45 55 60 62 63 44 46 Temnac 56 73 Bedard 72 Cove Bay 65 70 66 71 Nevidiskov 67 68 69 Bay Krasni 172°30’E Point 173°15’E 52°45’N 53°45’N

0 km 15 Contact U D Dashed where approximately located Fault Datum is mean sea level Dashed where fracture or fault is inferred Shoreline shown represents the approximate from mapping or aerial photographs; U, 45 line of mean high water upthrown side; D, downthrown side Strike and dip of beds 1971 Magnetic declination varies from 2° to 2°30’ EAST

Figure 1.2 (Continued):

Figure 1.2 (Continued): Map # Sample Name Age Map # Sample Name Age 1 AT45 38 SB80-5A 29.63±0.42 Ma 2 AT50 39 SB80-1A 3 AT16, 17 40 AT80-61A 34.93±0.4 Ma 4 AT6, 8, 11, 12 41 AT80-36 5 AT22 42 AT80-32 6.19±0.07 Ma AT24, 25, 27, 43 MB80-15 18.93±0.75 Ma 6 29, 134, 135 7 AT33 44 MB80-24 8 AT57 45 AT80-27, 27A 9 AT56 46 AT83 10 AT19 16.17±0.11 Ma 47 AT80-28 11 AT21 48 AT80-30A 12 AT3 49 AT114 13 AT53 50 AT100 14 AT144 51 AT101 15 AT140B, 143 52 AT80-81 16 AT61 53 AT80-79E 17 AT73 54 AT80-78 18 AT65 55 AT80-87 29.38±2.50 Ma 19 AT77, 78 56 AT80-89 20 HO9-9 57 AT80-76 40.28±0.12 Ma 21 HO9-2D 58 AT80-54 22 HO9-14B 59 AT80-55 23 HO9-22 60 AT80-8B 24 HO9-23C 61 AT80-4A 25 HO9-25B 62 AT80-1A 26 HO9-49B 63 CP9-8 27 HO9-42 64 CP9-16B, 16C 28 HO9-62A 65 AT80-46 29 CH9-18 66 AT80-48 30 CH9-16 67 AT80-43, 43A 31 CH9-11 35.57±2.58 Ma 68 AT80-22 34.59±0.99 Ma 32 CH9-8 69 AT80-19A, 20 33 CH9-43D 70 AT80-71 33.46±0.52 Ma 34 CH9-52A 71 MP80-10 35 SB80-20 72 MP80-8 36 SB80-18 73 MP80-2 37 SB80-6A

pyroclastics, and rounded beach material. The pillow lavas and flows of this series are several

meters thick; some of the flows have columnar jointing; and the mostly basaltic pillows are 0.5-6

m in diameter, but are mostly 1-2 m. The pyroclastic rocks in the basement series are nearly as thick as the pillows and flows (up to 30 m thick) and range from massive to sorted bedding.

Several fine-grained tuffs appear to have been deformed before they were consolidated,

indicating tectonic activity at the time of deposition. On Murder Point (Figure 1.2) and along the

northwestern shore, folded and faulted volcanic breccias are exposed. Evidence of deformation,

faulting, slumping, turbidity flows, crossbedding, and scour and fill features indicate a steep-

sloped submarine environment, possibly due to volcanic peaks, that underwent tectonic

movement. The fine-grained laminated sediments also indicate local quiet deep basin deposition

without currents or waves. Gates et al. (1971) interpreted the basement rocks to be the result of

submarine volcanism and erosion and placed their age as anywhere from Late Mesozoic to Early

Tertiary. More recently reported ages for the Basement Series include an amphibole K-Ar age of

41.5±2 Ma on an amphibolite (caused by an early metamorphic event that altered many of the earliest volcanic rocks) (Vallier et al., 1983; Pickthorn and Vallier, 1991), an 40Ar/39Ar age of

40.3±0.1 Ma (Kay et al., 2014), plagioclase and hornblende 40Ar/39Ar ages of 34.69±0.99 Ma

and 35.27±2.04 Ma, respectively, on a gabbro (Jicha et al., 2006), and a 31.9±1.4 Ma K-Ar age

on a quartz gabbro (Delong and McDowell, 1975).

The Basement Series is locally unconformably overlain by sedimentary rocks of the

Chunkisak, Nevidiskov, and Chirikof Formations (Gates et al., 1971). The Nevidiskov Formation

is less complexly deformed than the Basement series and is 300 m thick. This formation is composed mostly of conglomerates made from fragments of weathered basement with no bedding or evidence of turbidity currents. Gates et al. (1971) interpret this formation to be the

result of extensive erosion of the basement followed by uplift and deformation. They place the age of the Nevidiskov Formation as Middle Tertiary (Early Miocene to Late Oligocene). Based on the ~40.3 Ma age of the basement rocks and the 35 Ma age of the plutons (Table 1.2), which should be younger than this formation (based on cross-cutting relationships), we now know the

Chunkisak, Nevidiskov, and Chirikof Formations were likely deposited between 40-35 Ma.

The Chuniksak Formation unconformably overlies the basement series and conformably overlies the Nevidiskov Formation. It is composed of approximately a thousand meters of fine- grained laminated siliceous argillite and limy sedimentary rocks. This formation is highly faulted and deformed with evidence of rhythmic banding. The rocks are composed of volcanic ash and marine sediments and include ellipsoidal nodules 0.5-3.5 m in diameter. The warping of the beds around these nodules indicates that they formed as features on the seafloor rather than by the replacement of original material. The age of the diatom fossils found in this formation are somewhat ambiguous, but Gates et al. (1971) place its age as Early to Middle Miocene. They interpret the transition from the coarser bedding of the Nevidiskov Formation to the finer bedding of the Chuniksak Formation as reflecting a change to quieter marine deposition in a slowly subsiding basin with no large surrounding topographic features and a restricted connection to the open ocean.

The Chirikof Formation is less than 60 m thick, dips 55-85° S, has no exposed conformable contacts, and is cut by a diabase dike. It is composed of rounded-angular pebble- boulder conglomerate, coarse sandstone, carbonaceous shale and sandstone, and a lava flow that could be part of the basement. The main outcrops are on the eastern tip of the island (Chirikof

Point) near Matthews Mountain. Gates et al. (1971) interpret the Chirikof Formation to be the

result of subaerial deposition of materials by streams or waves along a boulder or gravel beach.

They place its age as Middle Tertiary.

The intrusive rocks on Attu can be separated into three categories: diabase and gabbro, albite granite, and hornblende and dacite porphyries. The diabases and gabbros make up approximately 2/3 of the intrusives and outcrop as NE-SW trending diabase sills and dikes in the northeast around Holtz Bay and Chichagof Harbor. The gabbros tend to form irregular bodies

scattered around the island that cut the Chuniksak and Chirikof Formations. Gates et al. (1971)

suggest they are older than the Miocene Massacre Bay Formation, and the new ages by Jicha in

Table 1.1 show that the plutons are 35-30 Ma.

Table 1.1: 40Ar/39Ar Ages for Attu and Kiska Islands (B. Jicha, Univ. of Wisconsin, personal communication) Description Location Age (Ma) Error 2σ Phase dated Attu Island AT80-32 Andesitic stock Matthews Mtn. 6.19 0.07 groundmass AT5 rhyolite Albite-Granites 16.17 0.11 groundmass AT19 rhyolite lava Albite-Granites 16.17 0.18 groundmass MB80-15 pillow basalt Massacre Bay 18.93 0.75 groundmass AT80-87 gabbro Krasni Point 29.38 2.50 plagioclase SB80-5A pillow basalt Sarana Bay 29.63 0.42 groundmass AT80-71* gabbro Krasni Point 33.46 0.52 plagioclase AT80-22* gabbro Krasni Point 34.69 0.99 plagioclase AT80-61A pillow basalt Jackass Pass 34.93 0.40 plagioclase CH9-11 Mg-andesite Chichagof H. 35.57 2.58 groundmass pillow lava AT80-76 basalt Krasni Point 40.28 0.12 plagioclase CH9-18 basaltic dike Chichagof 41.61 0.56 groundmass

Kiska Island KS-12-1 andesite 4.54 0.02 plagioclase KS-12-8 Mg-andesite 13.57 0.41 groundmass KS-12-9 dacite 15.29 0.11 groundmass KS-12-7 andesite 31.69 0.41 groundmass KS-12-13 basaltic andesite 31.84 0.70 groundmass KS-12-10 basalt 36.48 0.32 groundmass KS-12-20 rhyolite 38.54 0.65 plagioclase KS-12-12 basalt 39.09 0.49 groundmass TN-182.30.003 basaltic andesite 46.31 0.91 groundmass Note: Ages were analyzed by Brian Jicha at the University of Wisconsin using a MAP 215-50 mass spectrometer. See Appendix A for full method details. *Dates from Jicha et al., 2006.

Volumetrically minor rhyolitic volcanics and associated granitic intrusions form a nearly

parallel trend from the eastern side of Stellar Cove southwest to the northwestern side of

Abraham Bay. Yogodzinski et al. (1993) referred to these collectively as the Albite-Granite

Suite. These units are locally intruded by 1-2 m wide ferrobasalt and andesite dikes, which are

termed the Late-Dikes by Yogodzinski et al. (1993). The albite granite intrusions are surrounded by a large ~3.2 km contact alteration zone. Gates et al. (1971) suggest the Albite-Granites are post-Miocene, while new 40Ar/39Ar and zircon ages indicate that they are ~16 Ma (Table 1.1).

The hornblende and dacite porphyry stocks and sills occur all across Attu where they cut the

Chuniksak Formation, the basement, and the alteration halo of the albite granites. They have an

E-W alignment that parallels normal faults, some of which in southern and eastern Attu could be

post glacial in age. Near Stellar Cove, these faults trend N-S. Gates et al. (1971) place the

porphyries (quartz keratophyres and hornblende andesites and dacite porphyries throughout Attu)

as Late Tertiary to Early Pleistocene, possibly younger than the Albite-Granites or contemporaneous with the Massacre Bay Formation.

The Massacre Bay Formation unconformably overlies the Basement Series and can be

overlain by glacial deposits and Quaternary alluvium. This formation, which can reach upwards

of 150 m in thickness, is composed of conglomerate, coarse sandstone, tuffs, and basaltic to

andesitic lava flows. The poorly bedded basal conglomerate, which dips gently to the east, is

composed of rounded to subrounded basement fragments and angular basaltic andesite porphyry

fragments from the Massacre Bay Formation eruptions. The conglomerates and sandstones are

0.5-150 m thick, and the sandstone has bedding and scour structures. The lava flows are massive

with columnar jointing and contain hornblende, plagioclase, and augite. The coarse bedded volcanic breccia and tuff are ~23 m thick and are interbedded with the flows. There is also

evidence of volcanic mudflows or avalanches. Gates et al. (1971) place the deposition of the

Massacre Bay Formation after the intrusion of the gabbros, diabases, and Albite-Granites, and

the deformation and uplift of the basement, but before the glaciation. They interpret the flows, stream deposits, land avalanches and mudflows, and the presence of wood to indicate a subaerial, wooded environment where there was a step relief cut into the basement. We now know that the

Massacre Bay volcanics are ~19 Ma, older than the Albite-Granites and Late-Dikes, but younger

than the plutons and more northerly volcanics (Sarana Bay, Chichagof Harbor, and Holtz Bay)

(see Table 1.1).

The Faneto Formation, which unconformably overlies the Chuniksak Formation and the

basement, is greater than 455 m thick, has a moderate to gentle northward dip, and is cut by

normal faults. It is composed of coarse red subaerial clastics eroded from the Chuniksak

Formation, whose red color is due to the argillite eroded from the basaltic lava flows. The rounded-subangular pebble-boulder conglomerate has poor sorting and bedding, was deposited by streams, and fines upward. This formation also includes subrounded-angular coarse sandstones and graywackes and thin lenses of argillite. Gates et al. (1971) place the Faneto

Formation as younger than the diabase and albite granite intrusives, but older than the glaciation.

Its hornblende content places it as post-syn Massacre Bay Formation in the Late Tertiary to Early

Pleistocene. We now know that to be post-syn depositional with the Massacre Bay Formation it would have to be ~19 Ma, but no younger than 6 Ma (age of the calc-alkaline Matthews

Mountain series).

Gates et al. (1971) report that the structural features of Attu include normal faults, possibly with a large strike slip component, and broad and open folds sub-oriented to the faults.

The Chuniksak Formation dips to the north on the north shore of Holtz Bay and to the south on

the south shore, forming a shallow syncline. Overall, the entire island is a broad anticline with an

E-W trending axis, with smaller structures super imposed upon it. Faulting has been active

throughout the island’s history. Across the southern peninsula area the faults mainly trend E-W;

whereas across the western and northern parts of the island the faults mainly trend NE. There is

less faulting on the eastern third of the island. Gates et al. (1971) report faults on the western part

of the island that could still be active. There are also active faults on the eastern part of the

island. The faults on the island also line up with submarine canyons and lineaments. There is

evidence of a massive uplift occurring between the deposition of the Chuniksak and Massacre

Bay Formations in the Late Miocene to Early Pleistocene. The western 1/3 of the island

(separated by a line from Stellar Cove to Abraham Bay) has NE trending valleys, while the rest

of the island has ESE trending valleys.

The youngest rocks on Attu comprise the Calc-Alkaline Series and include the andesites

of Matthews Mountain, which form an eroded volcano on the eastern peninsula of the island.

They are compositionally similar to the volcanoclastics of the Massacre Bay Formation, which

were deposited on the eroded surface of the Basement Series on the opposite side of Massacre

Bay from Matthews Mountain. Previously reported ages for the Calc-Alkaline Series range from

15 to 5 Ma (Delong and McDowell, 1975; Vallier et al., 1983: Table 1). Rubenstone (1984) and

Yogodzinski et al. (1993) found that the basaltic rocks of the Basement Series are tholeiitic, and

have radiogenic εNd (+10.1 to 11.4), unradiogenic Pb, flat to LREE-depleted rare earth patterns,

and MORB-like Th/La ratios. The rhyolites and granites of the Albite-Granite Suite are similar

in character to differentiated rocks found on propagating rifts of mid-ocean ridges (e.g., Bonatti

et al., 1975; Christie and Sinton, 1981) with only slightly enriched LREE and slightly lower εNd

(+8.8 to 9.6). Yogodzinski et al. (1993) concluded that they were produced by open system

fractional crystallization of an Attu basaltic parent.

Rubenstone (1984) identified a subset of tholeiitic basalts with arc-like trace element

(e.g., Ta depletion) and isotopic signatures (high 207Pb/204Pb relative to 206Pb/204Pb), from the

southeastern part of the island, principally near Casco and Murder Points, which likely formed in

an arc-related environment. The Casco and Murder Point volcanics were later determined by

DeLong and MacDowell (1975) to be ~33 Ma.

The Attu Calc-Alkaline Series consists of andesites and dacites that are rich in large-ion- lithophile (LIL) and light rare earth elements (LREE), poor in high field strength elements

(HFSE) and are similar in composition to the calc-alkaline rocks that outcrop on the other Near

Islands (Yogodzinski et al., 1993). Although similar in some respects to calc-alkaline magmas of the eastern and central Aleutians, those of the western Aleutians are distinct in having lower Th and heavy rare earth element (HREE) concentrations and more depleted isotopic signatures.

Yogodzinski et al. (1993) proposed that they are produced by fractional crystallization of a Mg- andesite parent, similar to those of Piip Volcano to the west (just northeast of Bering Island in the Komandorsky Islands; Figure 1.1). The parent lava was proposed to result from melt- peridotite interaction in the uppermost mantle and low temperature fractionation in a compressive environment, which resulted in the eruption of smaller volumes of a more fractionated magma.

Yogodzinski et al. (1993) and Rubenstone (1984) came to several conclusions in their studies of Attu. Based on the presence of pillow structures, interbedded laminated marine sediments, hydrothermal sub-greenschist metamorphism, and MORB-like compositions, they concluded that magmatism on Attu began with the voluminous eruption of tholeiites produced

from a depleted peridotite NMORB source in an oceanic/backarc spreading environment. This progressed to a propagating oceanic rift environment (e.g., Bonatti et al., 1975; Christie and

Sinton, 1981), which produced the albite granites and rhyolites though open system fractional crystallization, perhaps accompanied by crustal and marine sediment assimilation (lowering εNd and increasing Pb isotopes). Activity culminated with the intrusion of the Late-Dikes.

Yogodzinski et al. (1993) interpreted the island arc tholeiites of the southeast as being derived from a mantle wedge contaminated by fluids or melts derived from the subducting slab saturated in a Ti mineral such as rutile, apparently contemporaneous with the earlier ocean ridge-like tholeiites. We show here that these early tholeiitic basalts largely have arc-like HFSE depletion and likely formed in the early stages of a tholeiitic island arc like the Marianas.

Yogodzinski et al. (1993), Vallier et al. (1994), and Scholl (2007) argued that because the

Pacific-North American plate motion has remained constant since ~43 Ma (Engebretson et al.,

1985), large volumes of early magmatism on the western Aleutian Ridge probably built the ridge before the shift to more oblique convergence at ~43 Ma. As we will show, the Attu Basement

Series is arc-related. In their interpretation, the shift from a transtensional to a transpressional tectonic environment around 15 Ma led to the eruption of the relatively small volume andesites of the Calc-Alkaline Series.

Background Geology of Kiska Island

Kiska is distinct from Attu in having historic volcanism and in that the older rocks have been significantly less studied. The oldest available ages for Kiska are the K-Ar age of 55.3 ±6.7

Ma from the Vega Bay Formation in Lief Cove on the southwestern coast of Kiska (Marvin and

Cole, 1978) and the 40Ar/39Ar age of 46±0.91 Ma for a Murray canyon dredge sample west of

Kiska (Jicha et al., 2006).

Coats et al. (1961) describes the 285 km2 Kiska Island (Figure 1.3) as being composed of

three general units that underwent various degrees of structural evolution. The oldest is the Vega

Bay Formation, which covers the southern half of the island, and whose age Coats et al. (1961)

suggested to be Middle Tertiary, based on similarities to Late Oligocene-Early Miocene fossil-

bearing units on Rat Island and Amchitka Island. In detail, the Vega Bay Fm dips to the west,

has a minimum thickness of 610 m, and is composed of pyroclastic rocks interbedded with

basaltic flows along with minor sandstone and conglomerate units composed of reworked

volcanic rocks. On the west side of the peninsula (south arm of Kiska Harbor) at the head of

Sargent Cove, pillow structures are common. The rocks on the southernmost tip of the island are

relatively fresh except for alteration of olivine and glass. Farther north in the central part of the

island, the Vega Bay volcanic rocks are heavily altered and contain few if any of the original

minerals. This heavy alteration is attributed to low grade metamorphism associated with the

intrusion of a gabbroic body west of Sargent Cove and a gabbroic body in the vicinity of the

southern coast of Gertrude Cove. Coats et al. (1961) considered the gabbroic plutons to be

Middle Tertiary in age and exhibit little alteration, mostly of olivine to iddingsite, and

plagioclase glomerocrysts that exhibit fracturing and alteration.

The Kiska Harbor Formation outcrops over 78 km2 in the north central part of the island.

This unit is considered to be Late Tertiary to Early Pleistocene in age as it unconformably

overlies the Vega Bay Fm and underlies the Pleistocene to historic flows from Kiska Volcano.

Coats et al. (1961) noted the resemblance of the Vega Bay Formation units to dissected

composite cones on Tanaga Island and Kanaga Island that are associated with Pliocene fossils.

Figure 1.3: A) Geologic Map of Kiska Island depicting geologic units, sample locations, and ages of dated samples. B) Legend of Geologic Map of Kiska. (after Coats et al., 1961 – plate 71)

Figure 1.3 (Continued):

The Kiska Harbor Fm is composed of subaerial lava flows, autoclastic breccia, pyroclastics, and

water-laid pumiceous sandstone and conglomerate composed of volcanic rocks. Lava flows are

predominant in the north and thin to the south; whereas water-laid sedimentary rocks are

predominant in the south, thin to the north, and interfinger in between. The flows are up to a hundred meters thick, vesicular, subaerial, exhibit columnar jointing, and have no pillow structures. The breccia is coarse, angular to subrounded basalt with a finer volcanic matrix that locally grades to massive flows. Coats et al. (1961) interpret this to mean that part of the flow erupted under water or that autobrecciation occurred as the flowing lava cooled. This formation is interpreted to be the remnants of an old composite volcano. Its eruptive source is thought to lie to the north, approximately where Kiska Volcano is now located.

Kiska Volcano is a composite cone located on the northernmost point of the island. It makes up the most recent volcanic unit and is assumed to be Late Pleistocene to historic in age

based on original flow textures and a lack of evidence of glaciation. Coats (1961) reported unconfirmed activity in 1905, 1943, 1947. The Smithsonian Global Volcanism Program reports eruptions in 1962, 1964, 1969, and 1990. The volcano is characterized by interbedded flows and pyroclastic units with flows up to 30 m thick that have red tops and bottom breccias in some places. The flows are porphyritic and are either hypersthene andesites with minor olivine or two pyroxene andesites. A small subsidiary cone truncated by marine erosion is located two miles south of Vulcan Point.

Three structural events are suggested by Coats et al. (1961), which generally coincide

with the depositions of the Vega Bay Formation, the Kiska Harbor Formation, and the eruptions

of Kiska Volcano.

1) The volcanic units of the Vega Bay Fm are the most deformed as they exhibit steep dips

and show evidence of folding. The N30°E and N60°W trending linear features are

interpreted to be related to normal faults. The submarine trenches surrounding Kiska

Island could be associated with these NE and NW trending faults. The westward dip of

the layers southeast of Gertrude Cove is interpreted as the limb of a large fold, which is

the dominant structure in this area. North of Gertrude Cove, the units dip away from the

gabbroic intrusion, indicating that the dominant structure in this area is a large dome.

2) The Kiska Harbor Fm units have dips of less than 15° that are likely related to the gentle

tilting of blocks between normal faults. At least two faults that run parallel to N60°W

trending lineaments have been down-dropped to create the lowlands south of Kiska

Volcano that separate the volcano from the older part of the island to the south. The

steeper dip of the underlying flows indicates that the structural changes were intermittent

during the eruption of this formation.

3) The dips of the units on Kiska Volcano are original to the deposition of the flows and

pyroclastic units, indicating that it has not undergone any significant structural activity.

In summary, the geologic history of Kiska Island started with the eruption of the basaltic flows and pyroclastics of the Vega Bay Fm from subaerial and submarine vents, occasional marine and subaerial erosion and redeposition of the volcanic material as conglomerates and greywacke, disturbance by block faulting, and the intrusion of small gabbroic bodies. The Vega

Bay and Kiska Harbor Formations are separated by an unconformity indicating an episode of erosion between the depositions of the two units. The next episode of development (the Kiska

Harbor Fm) was the growth of a composite cone to the north of the island formed by the Vega

Bay Fm. Fine dacite pumice filled valleys and sheltered areas, which was followed by breccia and flows of olivine basalt and pyroxene basaltic andesite. This was followed by a long period of erosion, uplift possibly due to block faulting, terrace cutting by wave activity, and Pleistocene glaciation. The last episode was localized volcanism on the northernmost tip of the island producing Kiska Volcano. Some of the earliest activity may have occurred during the Pleistocene glaciation and eruptions continued to historic times.

NEW STUDIES OF ATTU AND KISKA ISLANDS

Sample Collection and Analytical Methods

This study seeks to expand on previous studies of Attu’s complex history and to investigate Kiska’s history in order to place them both in the context of their locations on the

Aleutian Ridge and in the larger context of the Bering Sea Region. In order to accomplish this goal, it was necessary to augment the existing geochemical and geochronological data for Attu

Island and to acquire data for Kiska Island. To this end, new samples were collected and analyzed on Kiska, and additional analyses were made on Attu Island samples collected in previous studies.

Major elements on these samples were analyzed using inductively coupled plasma optical emission spectroscopy (ICP-OES) at Cornell University. Trace elements were obtained using inductively coupled plasma mass spectrometry (ICP-MS) at Cornell University and Colgate

University. Isotopic data was obtained using thermal ionization mass spectrometry (TIMS) at

Cornell University. For additional details on sample preparation and analysis see Appendix A.

Below is a description of these samples and their role in this study.

Figure 1.4: Chondrite normalized spider diagrams for Attu. Normalization values from McDonough and Sun (1995). A) 6 Ma, B) 16 Ma Late-Dikes, C) 16 Ma Albite-Granites, D) 19 and 30-40, E) 30-35 Ma, F) 30-35 Ma Plutons, and G) 40 Ma. (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984.)

Figure 1.4 (Continued):

Figure 1.5: Primitive mantle normalized spider diagrams for Attu. Normalization values from McDonough and Sun (1995). A) 6 Ma, B) 16 Ma Late-Dikes, C) 16 Ma Albite-Granites, D) 19 and 30-40, E) 30-35 Ma, F) 30-35 Ma Plutons, and G) 40 Ma. (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984.)

Figure 1.5 (Continued):

Figure 1.6: SiO2 vs. FeO/MgO Tholeiitic vs. Calc-Alkaline discrimination diagrams for Attu and Kiska. Black line delineates Tholeiitic/Calc-alkaline boundary (after Miyashiro, 1974). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Figure 1.7: SiO2 vs. K2O discrimination diagrams for Attu and Kiska. Black lines delineate low, medium, and high-potassium groupings (after Gill, 1981). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Attu Island Samples

The Attu samples for which there was sufficient information to place the sample into an age category were divided into 7 groups for plotting and interpretation based on their locations

(Table 1.2; Appendix B; Figure 1.2), ages (Table 1.1), and whether the sample was intrusive or extrusive. These divisions separate the different stages of magmatism more specifically then the

Formations mapped by Gates et al. (1971) and account for different types of magmatism that occur within the same general timeframe. Where possible, the ages of undated samples were estimated by comparing their stratigraphic relations, location, sample description, and chemistry to that of dated samples. Full sample descriptions and locations can be found in Appendix B. All of the samples in these groups were used in the plots. In addition, a few of these samples were chosen to plot on the chondrite and primitive mantle normalized spider diagrams (Figures 1.4,

1.5). These samples were chosen because they are representative of their groups. The following groupings are used.

Group 1: The 6 Ma group comprises the Calc-Alkaline Series (Figure 1.6), which is located on and around Matthews Mountain and includes an andesite stock (AT80-32) dated at

6.19±0.07 Ma (see Table 1.1) and two dikes of similar rock type and chemistry that are located on the north shore of Massacre Bay east of Alexai Point (AT83 and AT80-30A). This group has the youngest age on Attu Island as well as being distinct in chemistry from the older rocks.

Sample AT80-36 is a basaltic andesite collected near sample AT80-32 on Matthews Mountain.

Samples AT80-36 and AT80-30A are plotted on the spider diagrams (Figures 1.4A, 1.5A). New analyses conducted for this study include: major elements for AT80-36, and trace elements for

AT80-30A and AT80-36.

Group 2: Two groups described by Yogodzinski et al. (1993) are shown here to have ages

of approximately16 Ma. The first group of Yogodzinski et al. (1993) is the group of dikes that

intrudes the Albite-Granite series: the Late-Dikes. This group is referred to here as group 3 and is comprised of Yogodzinski’s Late-Dike group, which also includes one gabbroic intrusion. The

Late-Dikes group occurs along the shore of Steller Cove to the northwest of the band of rhyolites and albite granites (see map in Yogodzinski et al., 1993). These dikes are tholeiitic and range from basaltic to andesitic in composition (Figures 1.6, 1.7). They are incompatible element- enriched compared to the earlier ocean ridge-like pillows and breccias (Figures 1.4C, 1.5C). The

Late-Dike samples include the following samples: dikes AT8, AT11, AT16, AT21, AT22, AT27,

AT45, AT50, AT134, AT135, AT137, and AT138 and stock AT6. Samples AT16, AT21, AT27, and AT50 are plotted on the spider diagrams.

Group 3: The second group of Yogodzinski et al. (1993) is the Albite-Granite Series, which includes the NE trending rhyolite and albite granite suite. As shown in Table 1.1, the two rhyolites from the Rhyolite – Albite Granite series have yielded 40Ar/39Ar ground mass ages of

16.17±0.11 (rhyolite AT5) and 16.17±0.18 Ma (granite AT19). Other samples in this group are rhyolites (AT24, AT25, lava AT29, AT33, dike AT65, AT77, AT78, a sill in the Faneto Fm

AT143, and AT145) and albite-granites (stock AT3, stock AT12, AT17, AT53, stock AT56,

AT57, and AT140B). Samples AT3, AT5, AT19, AT29, AT56, and AT143 are plotted on the spider diagrams (Figures 1.4B, 1.5B). New analyses conducted for this study include: trace elements for AT19.

Group 4: The 19 Ma group includes samples from the north shore of Massacre Bay west of Alexai Point, which includes a brecciated lava (MB80-24), a pillow basalt (MB80-15) with a groundmass 40Ar/39Ar age of 18.93±0.75 Ma (Table 1.1), and a dike to the east of Alexai Point

(AT80-27A). These dikes are interpreted to be of similar age to the pillow basalts as they cut the

30-35 Ma pluton. They do not have the incompatible element-enriched chemistry of the 16 Ma

group (Figure 1.4D), or the calc-alkaline nature of the Matthews Mountain andesites (Figure

1.6), and are spatially closer to the Massacre Bay samples than to the Sarana Bay samples.

Sample MB80-15 is plotted on the spider diagrams (Figures 1.4D, 1.5D). New analyses

conducted for this study include: major and trace elements and isotopes for MB80-15.

Group 5: The 30-42 Ma group includes the Chichagof Harbor and Holtz Bay samples.

The Chichagof Harbor samples include: basaltic dikes (CH9-18, CH9-16), basaltic pillow lavas

(CH9-43D, CH9-52A), an andesitic pillow lava (CH9-8), and a dacitic pillow lava (CH9-11).

The Holtz Bay samples include: basaltic pillow lavas (HO9-14B, HO9-25B, HO9-22, HO9-42),

basaltic pillow breccias (HO9-23C, HO9-62A), a basaltic stock (HO9-49B), a basaltic andesite pillow breccia (HO9-2D), and a basaltic andesite pillow lava (HO9-9). These samples are very similar in type and chemistry. One Chichagof Harbor sample (CH9-11) has a groundmass

40Ar/39Ar age of 35.57±2.58 Ma (Table 1.1); and a basaltic dike in the sheeted dike complex on

the northern part of Fish Hook Ridge (CH9-18) has a groundmass 40Ar/39Ar age of 41.61±0.56

Ma. Thus, all of the Chichagof Harbor and Holtz Bay samples are assumed to be approximately within this age range. The Sarana Bay samples (basaltic flows and pillow lavas: SB80-5A, SB80-

6A, SB1A, SB80-20; dacitic pillow breccia: SB80-18,), one of which (SB80-5A) has a groundmass 40Ar/39Ar age of 29.63±0.42 Ma (Table 1.1), all have similar chemistry and are

assumed to be the same age. The Casco Point (basaltic andesite pillow breccias: CP8-16B, CP98,

CP916C) and Murder Point (basalt to basaltic andesite pillow breccias and pillow lavas: MP80-2,

MP10, and MP80-8) samples are assumed to be approximately 30 Ma based on the whole rock

K/Ar ages reported by DeLong and MacDowell (1975): one sample from Casco Point 28.3±1.2

Ma, one from just below Murder Point 33.46±.52 Ma, and one slightly inland of Murder Point

26.5±1.4 Ma. This group also includes the inland pillow lavas and massive diabases (basaltic pillow lavas: AT61, AT73; basaltic andesite massive diabase: AT100, AT114; basaltic massive diabase: AT144, AT101) that are most similar to the Chichagof Harbor and Holtz Bay samples in chemistry and do not share the 16 Ma enriched characteristics (Figure 1.4). Samples SB80-18,

CH9-11, CP916C, MP80-8, HO9-9, and AT73 are plotted on the spider diagrams. New analyses conducted for this study include: major elements for AT80-71, CH9-11, SB80-5A, SB80-6A, and SB80-18; trace elements for AT80-71, CH9-11, SB80-6A, and SB80-18; and isotopes for

CH9-11, CH9-18, CP8-16B, HO9-49B, and SB80-5A.

Group 6: The 30-35 Ma pluton group is composed of the northeastern pluton in Jackass

Pass near Matthews Mountain (gabbros: AT80-61A, AT80-65; diorites: AT80-28, AT80-27)

(AT80-61A; 34.93±0.4 Ma) and the gabbroic plutons on the Krasni Point Peninsula (AT80-22,

AT80-20, AT80-43A, AT80-43, AT80-48, AT80-54, AT80-71, AT80-78, AT8-87, and AT80-

89) (AT80-87; 29.38±2.5 Ma and AT80-22; 34.59±0.99 Ma), which is sometimes referred to as

Weston Mountain or the Aboud Plutons in earlier publications. These plutons form a distinct chemical group from the 30-35 Ma volcanic samples. Samples AT80-22, AT80-61A, and AT80-

87 are plotted on the spider diagrams. New analyses conducted for this study include: major elements for AT80-22, AT80-61A, AT80-87; trace elements for AT80-22, AT80-61A, and AT8-

87; and isotopes for AT80-22, AT8-87.

Group 7: The 40 Ma group includes the basaltic country rock that is intruded by the 30-

35 Ma plutons on Krasni Point. This group is separated from the 30-42 Ma group 5 volcanics because it is the host rock of the plutons. They also range to more depleted La/Sm and Sm/Yb ratios, are more enriched in K and Ba, and have more enriched εNd. One of these samples (basalt

AT80-76) has a plagioclase 40Ar/39Ar age of 40.28±0.12 Ma. This group includes the following

samples: AT80-1A, AT80-4A, AT80-8B, AT80-19A, AT80-46, AT80-55, AT80-76, AT80-79E,

and AT80-81. Sample AT80-76 is plotted on the spider diagrams. New analyses conducted for

this study include: major elements for AT80-76; trace elements for AT80-1A, AT80-4A, AT80-

8B, AT80-19A, AT80-46, AT80-55, AT80-76, AT80-79E, and AT80-81; and isotopes for AT80-

76.

Kiska Island Samples

The Kiska samples were divided into groups according to the ages of the dated samples

(Table 1.1) and the inferred ages of the undated samples that are based on comparisons of

location (Figure 1.3), sample characteristics, and chemistry to samples of known age. Unlike the

extensive work done on Attu to divide the island into specific geologic units, the generalized

geologic reconnaissance mapping on Kiska resulted in only three units, one of which is related to

the historic Kiska Volcano. In an effort to expand this description and investigate Kiska’s history

in more detail (especially the older and larger Vega Bay Formation), we use age divisions rather

than the previous mapped formation names below. For full sample descriptions and locations see

Table 1.2. All of the samples in these groups were used in the plots. In addition, all of the

samples were plotted on the chondrite and primitive mantle normalized spider diagrams (Figures

1.8, 1.9). With the exception of the analyses for samples Kiska62 and 54-445, all of the analyses

on the Kiska samples are new.

Group 1: The historic group includes two samples from Kiska Volcano on the northern

tip of the Island reported by George et al. (2003). These two samples are a basalt (55-445) and a basaltic andesite (Kiska62) that were collected just south of the summit of Kiska Volcano.

Group 2: The 5 Ma group includes two samples from the northern half of the island in the

Vega Harbor Formation, on the western side of North Head (northern arm of Kiska Harbor).

These samples were taken from a unit mapped by Coats et al. (USGS bulletin 1028R, 1961) as

Late Tertiary breccia and lava flows. One of these andesitic samples (KS-12-1) has an age of

4.54±0.02 Ma approximately the same age (K/Ar, 5.5±0.7 Ma) as the older composite cone that underlies the current Kiska Volcano (Marlow et al., 1973). These samples are fresh, unaltered andesites with large plagioclase phenocrysts (KS-12-1) and hornblende and plagioclase phenocrysts (KS-12-2).

Group 3: The 13-15 Ma group contains a 13.57±0.41 Ma dacite (KS-12-8) and a

15.29±0.11 Ma basaltic andesite (KS-12-9) from the eastern shore of the Vega Point peninsula on the southernmost part of Kiska and a seamount just off shore of Vega Point.

Group 4: The 31 Ma group contains a basalt (KS-12-5), a 31.69±0.41 Ma andesite (KS-

12-7), and a 31.84±0.7 Ma basaltic andesite (KS-12-13) from the southwestern shore of Kiska around the shore of Lief Cove.

Group 5: The 36.48±0.32 Ma basalt (KS-12-10) is from Vega Point.

Group 6: The 38-39 Ma contains 39.09±0.49 Ma basalts (KS-12-11, KS-12-12, KS-12-

15, KS-12-16), an andesite (KS-12-14), and a 38.54±0.65 Ma rhyolite (KS-12-20) from various locations around the southern part of Kiska (see Figure 1.3).

Group 7: The 46.31±0.91 Ma basaltic andesite (TN-182-30-003) is from Murray Canyon, a submarine canyon off the west coast of Kiska.

Figure 1.8: Chondrite normalized spider diagrams for Kiska. Normalization values from McDonough and Sun (1995). A) Historic and 5 Ma, B) 13-15 Ma and 31 Ma, C) 38-39 Ma, and D) 36, 38-39 Ma and 46 Ma. (Includes Kiska data from George et al., 2003.)

Figure 1.8 (Continued):

Figure 1.9: Primitive mantle normalized spider diagrams for Kiska. Normalization values from McDonough and Sun (1995). A) historic and 5 Ma, B) 13-15 Ma, and 31 Ma, C) 38-39 Ma, and D) 36, 38-39 Ma and 46 Ma. (Includes Kiska data from George et al., 2003.)

Figure 1.9 (Continued):

Table 1.2: Kiska Island Sample Locations and descriptions Sample Description Location Latitude Longitude KS-12-1 Andesite, western arm of N52° 0.487’ E 177° 35.376’ columnar North Head, jointing, North arm of plagioclase Kiska Harbor phenocrysts, little weathering KS-12-2 Andesite, western arm of N52° 0.487’ E 177° 35.376’ plagioclase and North Head, hornblende North arm of phenocrysts, Kiska Harbor slightly weathered KS-12-3 Basalt North of Lief N51° 56.926’ E177° 19.688’ Cove KS-12-4 Basalt, plag. and North of Lief N51° 56.794’ E177° 19.891’ cpx phenocrysts Cove KS-12-5 Basalt, plag. North of Lief N51° 56.608’ E177° 19.643’ phenocrysts, Cove mineral filled vesicles KS-12-6 Basalt breccia directly below N51° 56.608’ E177° 19.643’ KS-12-5, North of Lief Cove KS-12-7 Andesite North arm of N51° 56.249’ E177° 19.625’ Lief Cove KS-12-8 Andesite, small sea knob N51° 49.35’ E177° 19.401’ columnar, olivine off southern phenocrysts, sea Kiska water alteration

Table 1.2 (Continued): Sample Description Location Latitude Longitude KS-12-10 Basalt, ol. and West side of N51° 49.852’ E177° 18.046’ cpx. phenocrysts, Vega Point badly weathered, underlies breccia KS-12-11 Basalt, large West side of N51° 49.852’ E177° 18.046’ clast incased in Vega Point KS-12-10, mineral filled vesicles KS-12-12 Basalt, above west arm of Dark N51° 52.932’ E177° 13.024’ breccia unit, cpx Cove phenocrysts KS-12-13 Basaltic andesite, South arm of N51° 55.84’ E177° 18.633’ mineral filled Lief Cove vesicles KS-12-14 Andesite, plag. South of Lief N51° 55.408’ E177° 17.312’ and amph. Cove phenocrysts KS-12-15 Basalt, plag. N51° 55.078’ E177° 15.858’ phenocrysts KS-12-16 Basalt, ~15 m south of KS-12- N51° 54.526’ E177° 15.073’ wide dike in 15, possibly cuts Breccia, cpx, KS-12-15 plag. and ol. phenocrysts, fairly weathered KS-12-17 Gabbro, cpx and west arm of Jeff N51° 56.020’ E177° 28.870’ plag. phenocrysts Cove, 33m asl

Table 1.2 (continued): Sample Description Location Latitude Longitude KS-12-20 Rhyolite, center of N51° 56.108’ E177° 27.563’ columnar jointed, Gertrude Cove, very altered, south of stream, light green 2m asl matrix with altered plagioclase? TN-182-30-003 Basaltic andesite Murray Canyon dredge sample Note: see Appendix C for detailed thin section descriptions.

RESULTS

Major Elements

Major element concentrations for Attu and Kiska are listed in Tables 1.3 and 1.4, respectively, and in Appendix D. The Attu samples range from basalts to rhyolites with the most silicic rocks restricted to the period from 19 to 16 Ma. The pre-19 Ma rocks span the range from basaltic to andesitic with no systematic variation in SiO2 with time (Figures 1.6, 1.7). In contrast,

the Attu Calc-Alkaline Series is exclusively andesitic. Arc-related magmas are often classified as

either low-, medium-, or high-K based on their K2O content as a function of SiO2 (e.g., Gill,

1981), with K2O serving as a proxy for the general incompatible element enrichment of the

mantle source. The Attu samples plot primarily in the medium-K field on the plot of SiO2 vs.

K2O (Figure 1.7), although they scatter considerably. However, K is also a mobile element, and

Rubenstone (1984) and Yogodzinski et al. (1993) concluded that the variance in K2O content at a given silica level is partly the result of alteration; so it is unlikely that any true high-K rocks exist on Attu. This seems a reasonable conclusion as the high-K rocks are volcanic and are mostly older samples, and hence, they are more likely to have experienced alteration due to low grade metamorphism in the presence of sea water (Rubenstone, 1984; Yogodzinski et al., 1993). The thin sections of these high-K samples (AT50, AT80-27A, MP80-2, and HO9-49B) exhibit alteration characteristics, e.g., secondary chlorite, quartz, calcite, and opaques, that support this hypothesis. This was also the conclusion of Rubenstone (1984) based on a detailed study of their primary and alteration mineralogy. The Matthews Mountain samples (6 Ma), which are relatively fresh, are medium-K. The 30-35 Ma group is predominantly low-K with a few medium-K samples. As K2O does not correlate with εNd, the low-K samples of this group could be due to

larger extents of melting or alteration. Alteration is deemed the more likely cause based on the

similarity of their TiO2 content to other samples with the same SiO2 content

The Attu samples plot mostly within the tholeiitic field on the plot of SiO2 vs. FeO/MgO

with the notable exception of the 6 Ma Matthews Mountain samples, which plot within the calc-

alkaline field (Figure 1.6). Three other older samples, CH9-11 (35 Ma; point between Chichagof

Harbor and Holtz Bay), SB80-18 (30 Ma; Sarana Bay), and MB80-24 (19 Ma; Massacre Bay),

plot in the calc-alkaline field, but all three of these samples show severe alteration and/or vug

fillings and do not differ significantly in trace element chemistry from the tholeiites. Thus, their

apparent “calc-alkaline” nature may merely reflect this alteration due to the addition of

secondary minerals through alteration and precipitation in vugs, which when extensive enough,

could alter the MgO and FeO contents of a rock (Rubenstone, 1984).

Like the Attu samples, the Kiska samples range in composition from basaltic to rhyolitic;

however, there are fewer high silica samples, and both the mafic and silicic volcanism span the

history of the island. The Kiska samples are also predominantly medium-K, with the 38-39 Ma

age group having one sample each in the low-K and high-K fields. These deviations from

medium-K in the older volcanics are likely the result of low grade metamorphism as on Kiska,

rather than representing the presence of true low and high-K rocks. This is supported by the

degree of alteration observed in the thin sections (Appendix C). Basalt KS-12-15 (low-K) has chlorite present in its matrix and replacing plagioclase and serpentine replacing the olivine phenocrysts. In addition, the olivine phenocrysts are completely altered and the plagioclase phenocrysts are heavily altered with chlorite ingrowths. Rhyolite KS-12-20 (low-K) also has chlorite present in its matrix as well as other alteration products. It also contains extremely

altered plagioclase and fractured quartz phenocrysts both with some secondary fracture filling with chlorite.

The Kiska samples are similarly tholeiitic (Figure 1.6) with the exception of the historic and 5 Ma andesites from northern Kiska, the oldest 46 Ma dredge sample (TN-182.30.003) from

Murray Canyon, the 13 Ma sample from a small sea knob (KS-12-8), and the 38 Ma Gertrude

Cove rhyolite (KS-12-20). The latter three are severely altered, and like the Attu samples, their calc-alkaline signature may merely reflect alteration.

Table 1.3: Major Elements for Attu Island Samples AT80-22 AT80-36 AT80-61A AT80-71 AT80-76 AT8-87 Map # 68 41 40 70 57 55

SiO2 52.90 54.10 48.39 50.81 51.00 49.73 TiO2 2.02 0.64 0.70 0.87 0.94 0.85 Al2O3 13.18 17.42 19.05 17.48 14.50 20.10 FeO 16.31 6.97 9.83 11.35 13.57 10.67 MnO 0.28 0.16 0.15 0.19 0.28 0.15 MgO 2.99 5.98 6.68 5.31 6.54 3.98 CaO 7.04 8.00 10.80 10.42 5.76 10.82 Na2O 3.18 3.35 3.31 3.23 4.31 3.34 K2O 0.80 1.25 0.56 0.45 1.27 0.84 P2O5 0.20 0.15 0.11 0.11 0.14 0.10 Total 98.91 98.03 99.57 100.20 98.31 100.58

CH9-11 MB80-15 SB80-5A SB80-6A SB80-18 Map # 31 43 38 37 36

SiO2 64.83 49.15 49.66 50.14 65.76

TiO2 1.10 1.70 1.56 1.75 1.00

Al2O3 14.45 15.61 15.43 17.17 14.70 FeO 7.57 14.20 10.00 9.31 7.69 MnO 0.00 0.20 0.17 0.16 0.10 MgO 5.81 5.24 9.21 6.84 6.81 CaO 1.91 8.29 8.04 9.99 2.12

Na2O 4.00 3.68 3.54 4.41 4.08

K2O 0.25 0.13 0.61 0.28 0.08

P2O5 0.00 0.18 0.24 0.25 0.00 Total 99.92 98.40 98.47 100.28 102.34 All data analyzed at Cornell University by ICP-OES.

Table 1.4: Major Elements for Kiska Island Samples Kiska62† 54–445† KS-12-1 KS-12-2 KS-12-3 KS-12-4

SiO2 55.22 49.29 58.34 58.19 47.46 49.44

TiO2 0.75 0.82 0.69 0.65 0.84 0.74

Al2O3 18.68 15.64 17.87 17.13 17.13 20.91 FeO 7.88 10.06 7.36 7.14 12.13 10.68 MnO 0.15 0.15 0.14 0.13 0.27 0.18 MgO 3.8 10.7 3.43 3.46 6.55 5.21 CaO 8.47 9.93 7.44 7.46 9.99 9.89

Na2O 3.5 2.61 3.36 3.44 3.13 2.45

K2O 1.11 0.44 1.67 1.85 1.16 1.10

P2O5 0.17 0.11 0.13 0.13 0.09 0.09 Total 99.74 99.75 100.42 99.59 98.75 100.69

KS-12-5 KS-12-6 KS-12-7 KS-12-8 KS-12-9 KS-12-10

SiO2 49.72 51.56 58.42 56.48 67.91 50.66

TiO2 0.82 0.77 1.03 0.59 0.25 0.70

Al2O3 19.17 19.76 16.38 17.08 17.64 19.78 FeO 11.47 10.97 10.18 6.44 1.86 10.20 MnO 0.22 0.25 0.22 0.11 0.05 0.15 MgO 6.23 5.37 3.40 5.68 0.15 4.70 CaO 6.56 7.54 5.76 7.83 3.46 10.73

Na2O 5.72 4.18 4.21 3.85 5.18 2.54

K2O 0.57 0.57 1.37 1.08 1.80 0.39

P2O5 0.10 0.10 0.44 0.09 0.14 0.13 Total 100.59 101.07 101.41 99.23 98.43 99.98

KS-12-11 KS-12-12 KS-12-13 KS-12-14 KS-12-15 KS-12-16

SiO2 50.63 48.27 55.79 57.40 49.24 47.81

TiO2 0.73 0.78 1.08 0.87 0.83 0.67

Al2O3 18.82 19.20 16.34 15.95 18.79 19.00 FeO 10.81 11.06 10.61 8.05 11.54 11.31 MnO 0.18 0.18 0.23 0.17 0.19 0.18 MgO 5.99 6.31 3.71 2.46 5.41 5.70 CaO 11.58 12.13 4.88 6.10 8.20 11.69

Na2O 2.48 2.12 5.18 3.94 3.36 1.96

K2O 0.48 0.25 1.28 1.67 1.34 0.41

P2O5 0.13 0.08 0.44 0.29 0.09 0.09 Total 101.84 100.39 99.54 96.87 98.98 98.83

Table 1.4 (Continued): KS-12-17 KS-12-18 KS-12-20 TN-182.30.003

SiO2 52.25 50.00 74.97 55.95

TiO2 1.03 0.97 0.43 0.85

Al2O3 16.10 19.06 12.72 14.86 FeO 10.73 10.08 3.34 8.07 MnO 0.15 0.14 0.11 0.23 MgO 6.47 4.38 2.76 7.57 CaO 8.48 10.32 0.40 7.78

Na2O 4.01 4.50 3.13 5.17

K2O 1.16 0.61 1.31 0.97

P2O5 0.22 0.20 0.07 0.08 Total 100.60 100.26 99.24 101.52 † indicates data from George et al., 2003; all other data analyzed at Cornell University by ICP- OES.

Table 1.5: Trace Elements for Attu Island Samples AT19 AT80-1A AT80-4A AT80-8B AT80-19A AT80-22 AT80-30A Map # 10 62 61 60 69 68 48 Sc 3.79 37.06 32.30 45.28 9.42 44.59 13.06 V 0.68 389.7 282.3 483.2 9.51 384.3 131.7 Cr 102.6 25.89 10.00 27.89 2.21 4.74 34.26 Co 1.17 36.23 39.62 36.21 7.84 34.50 16.17 Ni 49.13 20.67 18.03 20.09 4.85 7.25 22.44 Cu 6.95 139.1 130.1 275.8 26.22 598.1 45.01 Zn 51.08 94.43 87.67 104.0 26.27 156.1 51.39 Rb 13.30 9.72 52.58 16.60 1.53 12.44 14.75 Sr 11.38 226.3 452.4 210.4 50.80 221.1 484.7 Y 43.58 19.37 18.99 24.24 47.54 41.12 12.54 Zr 756.9 60.99 57.64 19.80 41.45 42.68 99.99 Nb 23.18 0.65 0.54 0.53 1.28 1.05 1.51 Ba 214.9 131.1 927.2 192.7 21.40 197.0 277.7 La 20.41 4.67 4.76 2.68 6.63 3.80 7.57 Ce 58.14 11.47 11.50 7.02 17.68 10.25 18.10 Pr 7.42 1.91 1.87 1.21 3.05 1.87 2.55 Nd 31.32 9.46 9.32 6.57 15.46 10.21 10.66 Sm 7.53 2.76 2.68 2.37 4.87 3.58 2.37 Eu 1.80 0.93 1.16 0.90 1.22 1.27 0.83 Gd 7.72 3.09 3.11 3.03 5.86 4.59 2.51 Tb 1.37 0.54 0.52 0.59 1.11 0.89 0.38 Dy 8.42 3.36 3.25 3.92 7.21 6.02 2.09 Ho 1.76 0.72 0.69 0.86 1.54 1.34 0.43 Er 5.01 1.99 1.92 2.43 4.24 3.77 1.22 Tm 0.79 0.31 0.30 0.38 0.65 0.59 0.19 Yb 4.98 1.94 1.81 2.34 3.88 3.63 1.22 Lu 0.79 0.30 0.28 0.37 0.59 0.56 0.19 Hf 15.55 1.83 1.71 0.97 1.47 1.63 2.62 Ta 1.67 1.14 0.04 0.06 0.09 0.09 0.11 Pb 4.24 1.75 1.22 1.20 0.69 1.46 3.30 Th 2.50 0.39 0.41 0.16 0.41 0.23 1.09 U 1.29 0.18 0.23 0.05 0.16 0.10 0.64

Table 1.5 (Continued): AT80-36 AT80-46 AT80-55 AT80-61A AT80-76 AT80-79E Map # 41 65 59 40 57 53 Sc 20.24 40.67 45.47 33.44 47.73 35.36 V 168.4 395.6 494.1 277.1 397.0 273.2 Cr 163.2 60.74 26.71 169.7 31.00 64.40 Co 24.62 34.86 36.71 34.92 39.85 23.29 Ni 68.54 27.36 19.77 58.43 22.07 31.21 Cu 53.68 298.1 333.0 74.62 183.2 30.80 Zn 74.57 91.57 102.6 49.52 97.34 23.73 Rb 14.51 6.53 15.79 10.26 15.66 1.60 Sr 491.2 213.0 193.6 390.7 193.3 96.30 Y 14.26 18.44 24.65 13.86 24.44 21.11 Zr 86.86 13.37 18.67 24.02 64.45 37.51 Nb 1.25 0.40 0.52 0.28 0.68 0.56 Ba 325.4 141.3 189.7 80.25 401.3 24.17 La 7.32 2.33 2.72 2.70 5.01 4.56 Ce 17.69 5.96 7.12 6.79 12.58 10.26 Pr 2.67 1.00 1.22 1.15 2.11 1.67 Nd 11.69 5.44 6.73 6.01 10.53 8.20 Sm 2.75 1.87 2.36 1.85 3.09 2.42 Eu 0.93 0.74 0.92 0.63 1.10 0.75 Gd 2.85 2.42 2.99 2.12 3.51 2.86 Tb 0.44 0.47 0.59 0.38 0.61 0.52 Dy 2.45 3.11 3.91 2.37 3.86 3.30 Ho 0.51 0.67 0.88 0.50 0.83 0.72 Er 1.41 1.90 2.43 1.35 2.29 2.02 Tm 0.21 0.30 0.38 0.21 0.36 0.32 Yb 1.35 1.82 2.35 1.26 2.19 1.93 Lu 0.21 0.28 0.37 0.20 0.34 0.29 Hf 2.26 0.58 0.85 0.80 1.99 1.46 Ta 0.09 0.03 0.04 0.03 0.05 0.10 Pb 5.57 1.29 1.38 0.84 1.25 0.69 Th 0.90 0.12 0.18 0.20 0.45 0.34 U 0.52 0.05 0.06 0.11 0.19 0.15

Table 1.5 (Continued): AT80-81 AT8-87 CH9-11 MB80-15 SB80-6A SB80-18 Map # 52 55 31 43 37 36 Sc 45.71 37.74 19.34 35.86 33.30 20.67 V 522.9 477.5 237.9 407.4 255.8 203.9 Cr 19.88 22.93 6.54 2.63 210.48 10.25 Co 35.67 29.76 16.66 33.89 35.58 19.31 Ni 18.48 20.24 15.16 12.98 98.19 18.49 Cu 76.41 174.8 46.95 142.1 70.89 163.7 Zn 102.7 66.58 71.71 89.21 73.25 50.52 Rb 18.87 13.89 1.79 2.84 4.22 1.67 Sr 231.7 307.9 83.69 244.1 286.2 126.9 Y 25.84 16.80 29.61 24.41 38.16 29.39 Zr 25.15 25.47 77.13 51.93 157.9 71.99 Nb 0.55 0.36 2.03 0.83 7.43 2.31 Ba 167.2 143.0 34.67 49.66 51.85 31.48 La 3.42 1.97 3.52 2.64 6.59 5.27 Ce 7.90 4.94 10.22 7.64 18.93 12.53 Pr 1.30 0.83 1.80 1.34 3.05 1.99 Nd 6.93 4.58 9.63 7.46 14.80 10.17 Sm 2.40 1.63 3.03 2.62 4.21 3.05 Eu 0.94 0.70 1.14 1.10 1.45 1.09 Gd 3.08 2.15 3.75 3.21 4.88 3.82 Tb 0.59 0.42 0.69 0.62 0.89 0.69 Dy 3.93 2.81 4.44 4.03 5.55 4.37 Ho 0.87 0.62 0.96 0.87 1.19 0.94 Er 2.41 1.73 2.62 2.34 3.25 2.52 Tm 0.38 0.27 0.39 0.36 0.50 0.38 Yb 2.33 1.69 2.30 2.19 2.99 2.21 Lu 0.36 0.27 0.34 0.34 0.46 0.33 Hf 1.01 0.85 1.97 1.48 3.53 1.83 Ta 0.59 0.03 0.15 0.07 0.48 0.14 Pb 1.45 0.76 0.39 0.45 0.71 0.59 Th 0.17 0.14 0.19 0.11 0.43 0.21 U 0.08 0.07 0.25 0.07 0.22 0.15 Analyses performed at Colgate University by ICP-MS.

Table 1.6: Trace Elements for Kiska Island Samples Kiska62† 54–445† KS-12-1 KS-12-2 KS-12-3 KS-12-5 KS-12-6 Sc 19 34 13.88 15.79 36.79 28.85 28.95 V 191 279 200.9 202.5 411.0 328.8 299.6 Cr 11 554 62.76 73.26 65.43 61.22 62.42 Co 19.30 19.15 36.21 32.11 27.81 Ni 7 174 18.61 18.98 23.72 33.93 14.63 Cu 38 80 31.62 32.70 99.13 119.5 77.13 Zn 66 75 60.98 59.27 77.99 67.44 78.10 Ga 18.1 16.5 16.68 17.06 14.98 16.78 17.07 Rb 21.85 3.97 13.45 15.96 4.84 5.14 5.45 Sr 396 264 411.6 404.8 457.3 250.7 352.8 Y 21.69 15.2 18.08 18.04 15.52 15.77 17.30 Zr 102.2 63.2 102.8 110.5 36.40 39.63 40.62 Nb 2.35 1.18 3.60 4.07 0.74 0.49 0.76 Cs 0.53 0.26 0.10 0.08 0.22 Ba 331 135 314.9 360.4 115.6 105.8 215.9 La 8.85 4.1 7.52 7.86 1.93 1.88 2.24 Ce 20.59 10.64 18.04 19.07 5.48 5.50 6.18 Pr 3.02 1.7 2.62 2.73 0.99 0.98 1.10 Nd 13.19 8.13 10.69 10.93 4.75 4.82 5.23 Sm 3.52 2.52 2.93 2.97 1.77 1.73 1.86 Eu 1.12 0.85 0.85 0.86 0.58 0.58 0.60 Gd 3.5 2.69 3.26 3.34 2.51 2.53 2.66 Tb 0.56 0.43 0.50 0.51 0.40 0.41 0.43 Dy 3.49 2.62 2.95 2.91 2.43 2.44 2.62 Ho 0.73 0.53 0.59 0.60 0.50 0.51 0.55 Er 2.1 1.48 1.82 1.81 1.55 1.58 1.70 Tm 0.34 0.23 0.29 0.29 0.24 0.24 0.26 Yb 2.18 1.4 1.95 1.94 1.55 1.59 1.74 Lu 0.36 0.22 0.26 0.27 0.15 0.15 0.17 Hf 2.77 1.67 2.81 3.07 0.99 0.87 1.00 Ta 0.19 0.09 0.26 0.32 0.07 0.06 0.09 Pb 5.59 1.81 5.48 4.52 2.05 1.41 1.92 Th 0.936 0.231 1.78 1.68 0.16 0.13 0.21 U 1.936 0.423 0.90 1.07 0.12 0.08 0.09

Table 1.6 (Continued): KS-12-7* KS-12-8 KS-12-9* KS-12-10 KS-12-11 KS-12-12* KS-12-13* Sc 24.53 22.81 1.07 20.90 29.81 37.25 26.27 V 160.6 201.0 4.93 314.7 340.6 347.9 176.6 Cr 3.72 198.1 3.80 84.87 164.0 35.59 3.85 Co 19.00 22.21 1.76 28.49 35.21 39.64 19.76 Ni 6.66 72.94 2.65 31.60 31.58 35.60 7.24 Cu 42.50 50.92 3.20 73.26 107.0 128.3 53.86 Zn 116.3 55.74 33.63 80.16 74.21 74.39 132.6 Ga 16.95 18.08 17.97 Rb 18.96 12.99 33.67 3.93 4.16 3.54 17.45 Sr 400.5 351.6 385.9 569.0 596.2 346.7 428.2 Y 43.99 15.33 10.26 13.85 14.59 15.25 45.30 Zr 113.7 79.99 164.5 46.75 48.85 28.88 115.9 Nb 2.28 1.36 2.86 0.73 0.32 2.34 Cs 0.21 0.08 0.05 Ba 337.0 764.8 584.0 282.3 299.6 94.5 299.1 La 10.42 3.71 8.73 3.12 3.51 2.31 10.89 Ce 27.02 9.44 20.26 8.44 9.44 5.76 28.05 Pr 4.39 1.37 2.62 1.40 1.45 0.95 4.46 Nd 20.91 5.79 10.41 6.20 6.40 5.06 21.38 Sm 5.67 1.81 2.09 1.81 1.80 1.67 5.75 Eu 1.71 0.64 0.78 0.61 0.60 0.69 1.76 Gd 6.07 2.62 2.15 2.59 2.66 2.08 6.16 Tb 1.02 0.40 0.31 0.39 0.39 0.39 1.05 Dy 6.36 2.19 1.69 2.07 2.03 2.54 6.48 Ho 1.35 0.45 0.36 0.43 0.42 0.56 1.39 Er 3.85 1.39 1.07 1.36 1.33 1.54 3.92 Tm 0.59 0.22 0.18 0.21 0.20 0.25 0.62 Yb 3.63 1.44 1.19 1.39 1.34 1.49 3.76 Lu 0.57 0.18 0.20 0.14 0.13 0.23 0.59 Hf 3.43 2.07 4.14 1.19 1.18 0.91 3.47 Ta 0.14 0.13 0.23 0.13 0.07 0.03 0.77 Pb 6.69 4.52 8.50 2.95 2.57 1.28 5.45 Th 1.25 0.82 1.53 0.30 0.33 0.13 1.26 U 0.75 0.60 1.18 0.22 0.20 0.08 0.68

Table 1.6 (Continued): KS-12-20* KS-12-14* KS-12-15 KS-12-16* KS-12-17 TN-182.30.003* Sc 6.52 24.34 29.36 35.97 23.09 37.11 V 27.25 184.5 357.2 333.9 292.0 254.9 Cr 4.70 4.59 81.23 18.86 139.5 267.3 Co 4.37 17.06 30.03 40.47 29.34 39.73 Ni 3.68 7.98 16.64 28.75 48.41 75.46 Cu 5.71 56.40 142.7 214.6 82.55 70.71 Zn 46.31 94.20 75.76 73.93 70.95 126.9 Ga 17.86 16.80 Rb 21.64 16.87 8.84 7.37 12.73 10.36 Sr 41.24 387.2 1099.1 426.1 303.7 117.9 Y 37.23 40.03 13.92 15.56 20.46 19.76 Zr 200.0 130.9 39.52 40.60 69.28 47.40 Nb 3.40 3.14 0.70 0.54 3.94 0.54 Cs 0.16 0.19 Ba 393.9 277.7 440.9 128.3 235.6 96.99 La 9.55 9.38 1.83 3.50 5.39 2.05 Ce 21.64 23.83 5.88 8.31 15.89 5.58 Pr 3.77 3.82 0.98 1.33 2.18 0.92 Nd 16.76 17.89 4.70 6.57 9.65 5.05 Sm 4.42 4.79 1.64 1.94 2.84 1.81 Eu 1.08 1.39 0.65 0.71 0.79 0.74 Gd 4.77 5.23 2.56 2.28 3.53 2.44 Tb 0.85 0.90 0.39 0.41 0.57 0.48 Dy 5.47 5.63 2.15 2.62 3.29 3.19 Ho 1.23 1.22 0.45 0.56 0.69 0.70 Er 3.67 3.44 1.41 1.56 2.06 1.93 Tm 0.62 0.55 0.21 0.25 0.31 0.31 Yb 4.02 3.42 1.41 1.50 2.00 1.86 Lu 0.66 0.53 0.12 0.23 0.23 0.28 Hf 5.94 3.79 0.81 1.19 1.88 1.35 Ta 0.25 0.20 0.10 0.03 0.27 0.03 Pb 4.22 6.97 2.45 1.84 2.04 0.58 Th 2.47 1.50 0.16 0.38 0.72 0.15 U 1.34 0.82 0.07 0.21 0.41 0.15 † indicates data from George et al., 2003; * indicates analyses performed at Colgate University; all other samples were analyzed at Cornell University.

Trace Elements

Trace element concentrations for Attu and Kiska are listed in Tables 1.5 and 1.6, respectively, and in Appendix D. Attu samples with ages between 16 and 40 Ma generally show similar levels of rare earth element (REE) enrichment and have flat to slightly depleted light rare earth element (LREE) patterns, although some exhibit slight LREE enrichment (Figure 1.4).

Some of the most silicic samples have negative Eu anomalies, resulting from plagioclase fractionation. The 40 Ma sample, AT80-76, is slightly LREE enriched. The 16 Ma Albite-

Granite Series samples (rhyolites: AT24, AT25, AT29, AT33, AT65, AT77, AT78, AT143, and

AT145; and albite-granites: AT3, AT12, AT17, AT53, AT56, AT57, and AT140B) have enriched LREE, are the most enriched overall in REE, and have negative Eu anomalies. The 16

Ma Late-Dikes (AT11, AT50, AT16, AT27, AT22, AT45, AT138, AT8, AT135, AT137,

AT134, AT21, and AT6) have slightly enriched LREE, small negative Eu anomalies, and are more enriched overall than other samples with similar silica contents (Figure 1.4). Their enrichments compared to the other groups combined with their location and presumed age suggests that they are related to the Albite-Granites (Yogodzinski et al., 1993). The Matthews

Mountain samples have very enriched LREE patterns with the lowest concentration of heavy rare earth elements (HREE), but nonetheless, they have lower LREE concentrations than the Albite-

Granites. The 6 Ma samples exhibit a concave HREE pattern that is characteristic of amphibole.

The Kiska samples have the same range of REE concentrations as the Attu samples

(Figure 1.8). Kiska samples older than 5 Ma have primarily flat to slightly depleted LREE with a few slightly enriched LREE higher silica samples. The Middle Miocene dacite, KS-12-9, is very

LREE enriched and exhibits an amphibole fractionation pattern. Some of the more silica rich samples have slight negative Eu anomalies. The historic and 5 Ma samples are also LREE

enriched, similar to the Matthews Mountain samples on Attu; however, the 5 Ma Kiska samples have negative Eu anomalies, which are not seen in the historic Kiska or Matthews Mountain samples.

As Figures 1.5 and 1.9 show, the majority of the older samples on both Attu and Kiska exhibit some of the incompatible element enrichment patterns common in arc-related settings, with varying degrees of depletion in Th, Nb, and Ta and varying degrees of enrichments in Ba,

U, K, Pb, and Sr. It should be noted that especially for the older samples that have undergone sea water alteration and low-grade metamorphism, some of these elements (Ba, K, and Sr) are likely to have been altered or metamorphosed.

On Attu, the ~ 6 Ma Matthews Mountain calc-alkaline andesites exemplify the typical

HFSE depletion of subduction arc patterns. However, other Attu samples can differ in incompatible element patterns. The 19 Ma Attu samples have Ba, U, K, and Sr enrichments and depletions in Th, Nb, and Ta; however, Pb is depleted. This Pb depletion is typical of MORB, but is not common in arc-related volcanics. The 30-40 Ma samples have smaller Ba, U, and K enrichments and Pb and Sr depletions. The 30-40 Ma samples can exhibit either enrichment or depletion in Sr and have Pb depletions. The 30-35 Ma plutons (Figure 1.5) are enriched in Sr, have only small enrichments in Pb, but are the most depleted in Nb and Ta. The 40 Ma country rock sample is enriched in Ba and K, slightly enriched in Sr and U, shows no Pb enrichment, strong Nb and Ta depletions, and a small Th depletion. It is notable that the LREEs of the oldest samples are more depleted than those of most Aleutian volcanic rocks (see Yogodzinski et al.,

1993); however, this LREE depletion is also observed in some lavas from western Pacific arcs like the Mariana Arc (e.g., Reagan et al., 2010; Stern et al., 2003). The 16 Ma albite-granites and rhyolites are distinct because they lack the Ta depletion commonly seen in arc settings. The Attu

16 Ma Albite-Granite Series samples have very small relative enrichments in Ba, U, and K and small depletions in Th, Ta, and Sr. The 16 Ma basaltic to andesitic Late-Dikes exhibit similar

trends to the Albite-Granites. However, they have slightly larger enrichments and depletions in

these elements and lower overall concentrations, though they are still higher than other basaltic

and andesitic samples.

Unlike the somewhat variable patterns observed on Attu, the Kiska samples

systematically display typical arc-related incompatible trace element patterns, including alkali

and alkaline earth element enrichment, Ta and Nb depletion, and Pb enrichment, with a couple of

exceptions discussed below.

The Gertrude Cove rhyolite (38.54 Ma, KS-12-20) is exceptional in its Sr depletion

(Figure 1.9D). This sample has a negative Eu anomaly, and we attribute the Sr depletion to

plagioclase fractionation. It has intermediate Ba/La and the highest Th/La among the Kiska samples, although the 5 Ma andesites have only slightly lower Th/La. Alteration likely affected the Ba concentrations or the Gertrude Cove rhyolite.

A second exceptional sample is the 46 Ma basaltic andesite (TN-182) dredged from

Murray Canyon, which exhibits extreme Nb-Ta depletion, is LREE depleted (lowest Sm/Yb and

La/Sm), and lacks Ba and Pb enrichments (Figure 1.9D). This sample has been modified by sea

water alteration and low-grade metamorphism.

The Kiska samples with higher silica have higher overall incompatible element

concentrations (Figure 1.10). Some of the higher silica samples have negative Eu anomalies

(Figure 1.8), indicating fractional crystallization of plagioclase. The historic and 5 Ma Kiska samples look much like the 6 Ma Matthews Mountain samples erupted around the same time.

Figure 1.10: La/Sm vs. SiO2 for Attu and Kiska. La/ Sm normalized to C1 chondrite (ratio divided by 1.62) (McDonough and Sun, 1995). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Only the 31 Ma Kiska samples (KS-12-5, KS-12-7, KS-12-13; Figure 1.9B) have Zr and Hf depletions (Figure 1.9B), which is due to incomplete dissolution during sample preparation.

Due to the altered nature of the older rocks, especially on Attu, Ba, Rb, K, Pb, and to a lesser extent U and Sr could have been affected by alteration. Thus, anomalies in the abundances of these elements, or lack thereof, may not be diagnostic of their tectonic setting. Nb, Ta, and Th are much less susceptible to alteration influences and are thus likely the original magmatic concentrations. The subsequent discussion will focus mainly on these elements in combination with the rare earths.

Trace Element Ratios of Attu and Kiska Islands Compared to North Pacific MORB and Adak

Island in the Central Aleutians

When examining the trace element data it is helpful to compare it to both North Pacific

MORB and magmatic rocks from Adak Island in the central Aleutian Arc, east of Bowers Ridge

and west of the Alaska Peninsula (Figure 1.1). Adak Island has magmatic rocks ranging from 38

Ma to ~100 Ka in age (e.g., Kay and Kay, 1994), making it one of the few islands that includes

rocks from all stages of the arc’s history. Adak has been extensively studied, and there is a large

amount of data available for both the older plutonic (Citron et al., 1980; Kay, 1983; Kay et al.,

1983; Kay et al., 1990) and younger volcanic rocks, which makes Adak an ideal island to use in

comparing the central Aleutians to the western parts of the Aleutian Arc. Many other studies

have used Adak as a comparison due to these factors (e.g., Perfit et al., 1980; Conrad and Kay,

1984; Kay and Kay, 1985; Kay et al., 1990). The Adak field plotted in this study includes analyses from the Eocene Finger Bay Volcanics (Rubenstone, 1984) and the Holocene Mount

Moffett and Mount Adagdak volcanoes (modern volcanics).

Kiska samples older than 5 Ma have slightly depleted to slightly enriched LREE (see

La/Sm ratios in Figures 1.11, 1.10) with a similar range of concentrations to the Attu samples

(excluding the 16 Ma Albite-Granites and Late-Dikes). La/Sm and Sm/Yb both increase with time with the oldest samples plotting in the MORB field. In Sm/Yb-La/Sm space, the older

Kiska and Attu mafic samples overlap the North Pacific MORB field considerably (Figure 1.11),

but they also define a trend toward the Adak field (Plutons and Finger Bay Volcanics), and a few samples, most notably the younger calc-alkaline ones, plot within the Adak field. On Attu, the 16

Ma Albite-Granite Series defines a separate trend of increasing La/Sm with decreasing Sm/Yb.

There is a trend of more LREE enrichment (i.e., higher La/Sm) with higher silica content (Figure

1.10) for both Kiska and Attu. Like Attu, Kiska does not exhibit a trend in Sm/Yb with SiO2, but

it does show an increase in La/Sm with SiO2.

Nb and Ta depletion is a common characteristic, perhaps the defining one, of arc-related volcanism (e.g., Gill, 1981). We use the Ta/La ratio as a measure of that depletion. MORB have average Ta/La ratios of 0.062±0.021 (Hickey et al., 1986; White and Klein, 2014). Consequently,

we group samples into arc-like and ocean ridge-like on the basis of Ta/La ratios less than or

greater than ~0.04, respectively. On this basis, both arc-like and ocean ridge-like samples occur

on Attu (Figure 1.12). The ocean ridge-like category includes the 16 Ma Albite-Granite Series

and the Late-Dikes, the 19 Ma Massacre Bay samples, most of the 30-35 Ma volcanic samples,

and some of the 30-35 Ma pluton samples. The arc-like samples include Matthews Mountain (6

Ma), the 40 Ma sample, the samples from Casco Point and Murder Point (30 Ma), and 30-35 Ma

pluton samples AT80-61A (NE pluton), AT80-22 (SE pluton), and AT80-87 (SE pluton). Kiska

is predominantly arc-like with the exception of the 39 Ma KS-12-15. It should also be noted that

due to the powdering technique previously used on previous samples collected on Attu, some of

the previously ground samples, and although those with known contamination were not

considered, there is still the possibility for some Ta contamination. This could result in the Ta

values being too high. Thus, this contamination should be kept in mind when interpreting the Ta

values and the degree of Ta depletion.

The Th/La ratios of the older Attu samples are mostly ocean ridge-like (Figure 1.13),

trending into the lower half of the Adak field. The Albite-Granites do not follow this trend, but

this is a consequence of fractional crystallization due to the samples more evolved nature as the

Late-Dikes from the same age and area are MORB-like (Yogodzinski et al., 1993). The

Matthews Mountain, Murder Point, and the pluton samples from the north shore of Massacre

Bay also have Th/La greater than MORB. Overall, the Kiska Th/La range is shifted to slightly

higher Th/La than the Attu range and falls predominantly in the Adak field.

The Ba/La ratios scatter considerably, particularly for the older Attu samples, ranging from ocean ridge-like to values well above those of the central and eastern Aleutian Islands, such

as Adak (Figure 1.13). This is likely a consequence of the mobility of Ba, which can be either

added, or perhaps lost in some cases, from the rock during the low-grade metamorphism

apparent in many of the older samples. This is supported by the observation that the youngest,

least altered samples have Ba/La ratios similar to those of Adak and other Aleutian Islands.

Figure 1.11: Sm/Yb vs. La/Sm normalized to C1 chondrite (ratios divided by 0.92 and 1.62, respectively) (McDonough and Sun, 1995) for Attu and Kiska. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org) and Adak field in yellow (Kay and Kay, 1994; unpublished Cornell University database). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Figure 1.12: Ta/La vs. Ba/La for Attu and Kiska. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org) and Adak field in yellow (Kay and Kay, 1994; unpublished Cornell University database). Black line divides low Ta/La arc-like samples from high Ta/La ocean ridge-like samples. (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Figure 1.13: Th/La vs. Ba/La for Attu and Kiska. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org) and Adak field in yellow (Kay and Kay, 1994; unpublished Cornell University database). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

87 86 Figure 1.14: Sr/ Sr vs. εNd for Attu and Kiska. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org) and Adak field in yellow (Kay and Kay, 1994; unpublished Cornell University database). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Isotope Ratios

Sr and Nd isotopes are reported in Table 1.7 and Appendix D and are plotted in Figure

1.14. The Nd isotope ratios of Attu are higher than most other values observed in the Aleutian

Arc and island arcs globally (e.g., Rubenstone, 1984; Yogodzinski et al., 1993, 2015; Kay and

Kay, 1994; Kay et al., 1986), and they largely overlap the range of North Pacific MORB. Sample

from Kiska Island have lower εNd values more typical of the Aleutian Arc (e.g., Adak).

87 86 Unlike εNd, Sr/ Sr ratios differ from MORB values and are systematically higher,

ranging to greater than 0.705. While it is not uncommon for island arc volcanics to plot to the

high 87Sr/86Sr side of the Sr-Nd isotope mantle array, many of these 87Sr/86Sr values are high

even for island arcs, and they are generally higher than those observed for the younger volcanics

of the Aleutian Arc. In addition, there is a trend to lower 87Sr/86Sr with time. The most likely

explanation for this is that as was originally pointed out by Rubenstone (1984) and Yogodzinski et al. (1993), the Attu samples experienced hydrothermal alteration by seawater during low grade

metamorphism, which increased the 87Sr/86Sr values of the older samples.

The Attu samples, excluding the 16 Ma Albite-Granite Series, which form their own trend in La/Sm-Sm/Yb space (Figure 1.11), exhibit a broad, scattered anti-correlation of La/Sm with εNd (Figure 1.15), ranging from ocean ridge-like values to values more typical of the

Aleutian Arc as exemplified by the Adak samples. The Kiska samples show a similar trend, but

ranging to lower εNd and higher La/Sm than the Adak samples. This trend is largely controlled by

the higher La/Sm and lower εNd of the younger samples, but the older samples also show a small

trend in and of themselves.

The Attu samples generally plot within the same range on a plot of Th/La versus εNd;

however, the Albite-Granites have higher Th/La on average at the same εNd (Figure 1.16).

Notably, however, those samples with ocean ridge-like εNd also have ocean ridge-like Th/La.

Kiska forms a trend of increasing Th/La with decreasing εNd. This trend, like that of La/Sm with

εNd, falls between the MORB and Adak fields with a few of the younger samples extending beyond the Adak field.

The Sr isotope ratios of the Kiska samples (Figure 1.14) range from Adak-like to higher values in the older samples, with the higher values almost certainly due to hydrothermal alteration. εNd values (Figure 1.14) are intermediate between those of the Attu and Adak samples with the older samples falling in the lower end of the Attu range and the upper end of the Adak range.

Figure 1.15: La/Sm normalized to C1 chondrite (ratio divided by 1.62) (McDonough and Sun, 1995) vs. εNd for Attu and Kiska. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org) and Adak field in yellow (Kay and Kay, 1994; unpublished Cornell University database). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Figure 1.16: εNd vs. Th/La for Attu and Kiska. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org) and Adak field in yellow (Kay and Kay, 1994; unpublished Cornell University database). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

Table 1.7: Sr and Nd Isotopes for Kiska and Attu Islands 87Sr/86Sr ±2σ 87Sr/86Sr* 143Nd/144Nd ±2σ ε ±2σ ε * Nd Nd Kiska Island

KS-12-3 0.703376 13 0.703376 0.513133 16 9.65 0.32 9.65 KS-12-4 0.704203 7 0.704203 0.513056 20 8.14 0.38 8.14 KS-12-5 0.703858 8 0.703831 0.513128 18 9.56 0.36 8.63 KS-12-6 0.703747 10 0.703747 0.513118 12 9.36 0.24 9.36 KS-12-7 0.705245 8 0.705190 0.513094 11 8.89 0.22 8.22 KS-12-8 0.703191 14 0.703170 0.513062 11 8.27 0.22 7.95 KS-12-9 0.703177 6 0.703128 0.512985 14 6.78 0.28 6.53 KS-12-10 0.703077 14 0.703067 0.513085 11 8.72 0.22 7.91 KS-12-12 0.703230 6 0.703217 0.513132 22 9.64 0.42 8.62 KS-12-13 0.704556 10 0.704509 0.513118 13 9.37 0.26 8.69 KS-12-14 0.703465 13 0.703403 0.513105 14 9.10 0.28 8.27 KS-12-15 0.704307 13 0.704294 0.513117 16 9.34 0.30 8.29 KS-12-16 0.703349 10 0.703326 0.513124 17 9.48 0.34 8.59 KS-12-18 0.703232 8 0.703232 0.513115 13 9.30 0.26 9.30 KS-12-20 0.705297 11 0.704611 0.513117 16 9.34 0.30 8.54 TN-182.30.003 0.704619 8 0.704619 0.513147 17 9.93 0.34 9.93 Attu Island AT80-22 0.703129 6 0.703129 0.513121 16 9.42 0.30 9.42 AT80-71 0.703581 13 0.703554 0.513151 11 9.10 0.22 8.71 AT80-76 0.704219 7 0.704082 0.513074 20 8.51 0.38 7.53 AT8-87 0.703651 8 0.703615 0.513119 14 9.39 0.28 8.73 CH9-11 0.704059 8 0.704039 0.513180 11 10.58 0.22 9.70 CH9-18 0.703413 14 0.703388 0.513185 8 10.67 0.16 9.83 CP8-16B 0.703381 7 0.703341 0.513119 14 9.39 0.28 8.44 HO9-49B 0.703940 8 0.703755 0.513190 12 10.77 0.24 9.97 MB80-15 0.703350 8 0.703350 0.513201 13 10.98 0.26 10.98 SB80-5A 0.703276 13 0.703148 0.513159 16 10.16 0.32 9.48 * indicates age corrected data.

DISCUSSION

Magmatic and Tectonic Evolution of Attu and Kiska Islands

Attu Island

Yogodzinski et al. (1993), using a smaller data set, grouped all of the volcanics and plutons, which we now know range in age from 41 to 16 Ma into a single Basement Series, excepting only the Albite-Granite/Late-Dikes and the gabbros associated with the Late-Dikes.

They categorized the Casco Point and Murder Point volcanics as arc-like and the rest of the volcanics in his data set as MORB-like. No geochemical data were available for Massacre Bay, the Krasni Point plutons, the country rock of the plutons dated at 40.28±0.56 Ma, or the Jackass

Pass plutons and the 40.61±0.56 Ma age of the sheeted dike Fish Hook Ridge complex was unknown. In this paper, we use our extended geochemical and geochronological data set to investigate the tectonic environment in which the volcanics and plutons of Attu were emplaced.

We can then compare samples erupted at different times to determine how the magmatic environment on Attu changed over time. We then interpret Attu’s magmatic and tectonic evolution into the larger context of the Bering Sea Region. Based on this, we are able to significantly revise the evolution of Attu inferred in previous studies (Yogodzinski et al., 1993;

Shelton, 1986; Rubenstone, 1984; Vallier et al., 1994).

Our data is consistent with these earlier studies, which found that the evolution of Attu began by 41 Ma with the voluminous submarine eruption of tholeiitic basalts. These were subsequently intruded by the 30-35 Ma dikes and plutons. However, whereas Yogodzinski et al.

(1993) concluded that the earliest magmas were derived from a mantle similar to the MORB- source, our oldest dated Attu sample (41.1 Ma, AT80-76) is distinctly different from MORB in its Sr and Nd (εNd = 7.5) isotopic compositions (Figure 1.14). Furthermore, this sample and the

other earlier pillow lavas and volcanoclastics have flat to only slightly LREE depleted rare earth patterns, such as are found in the IBM and Japan (Reagan et al., 2010; Stern et al., 2003; Abe et al., 1998; Shuto et al., 2015), which are characteristics of arc tholeiites. They also have little to no Pb and Sr enrichments (Figures 1.4, 1.5) and Ta/La and Ba/La ratios (Figure 1.12) that are distinctly arc-like, though the Ba/La ratios could be due to alteration. While the Ba/La and

87 86 Sr/ Sr values were likely affected by hydrothermal alteration, the REE, Ta/La, and εNd values

are not easily explained this way. Hence, the overall picture is clear; the earliest magmatic rocks

preserved on Attu were generated above a subduction zone that had not experienced much

continental sediment and/or fluid input from previous subduction. It should be noted that the flat

to depleted LREE patterns seen in all but the calc-alkaline 6 Ma Matthews Mountain samples on

Attu are not seen east of Kiska Island in the Aleutian Arc. However, similar LREE-depleted

patterns are seen in the Mariana Arc (Reagan et al., 2010; Stern et al., 2003). This similarity will

be discussed further in Chapter 3. Their geochemical signature along with the large volume magmatism that occurred at this time suggests a more classic orthogonal subduction environment, with an initial fast build-up of the Aleutian Ridge as has been inferred by others

(e.g., Scholl et al., 1987; Kay et al., 1990; Jicha et al., 2006).

Arculus et al. (2015) describes this type of subduction initiation as induced initiation, which is caused by external forces such as ridge push. In the case of the Aleutian Arc, this would be the compressive force caused by the opposing motions of the extrusion of blocks of material from the Beringian Margin and the motion of the Pacific Plate creating stress that was relieved by the beginning of subduction along the Aleutian Arc (Scholl, 2007). The high melt volume could be due to a large influx of water when subduction initiated.

This induced subduction environment results in lower Th/La and La/Sm ratios (as larger

volumes result in lower incompatible element concentrations) and a slab influence contaminating the mantle wedge, resulting in a MORB signature with a slight overprinting of the strongest subduction signatures. This indicates that subduction was developed enough by 40 Ma to have already influenced the mantle wedge, but only minimally. Alternatively, this signature could be

characteristic of the mantle under the western Aleutians, with the mantle under the central

Aleutians having a different character (Yogodzinski et al., 2015; Yogodzinski et al., 2017; Kay et

al., 1986). This could be due to differences in the previous tectonic configuration of the area, i.e.,

having Bowers Ridge and the Kamchatka subduction zone nearby could lead to a different

remnant signature in the mantle than is present on the eastern side of the Bering Sea region.

From ~41-35 Ma the Chichagof Harbor (CH samples) (CH9-11: 35.57±2.58 Ma; CH9-

18: 41.61±0.56 Ma), Holtz Bay (HO samples), and inland samples span the range of MORB-like

Ta/La ratios, but these high Ta/La ratios are likely due to Ta contamination during sample

processing (Figure 1.12). They also have less enriched εNd and Th/La (Figures 1.13, 1.14, 1.16),

depleted to arc-like Ba/La (Figure 1.12), and depleted to flat LREE with REE concentrations

slightly higher than those of the plutons (Figure 1.4), but they also have large Ta, Nb, and Th

depletions as well as Ba, U, and K enrichments (Figure 1.5). The inland samples (gabbro: AT80-

71; basaltic pillow lavas: AT61, AT73; basaltic andesite massive diabase: AT100, AT114;

basaltic massive diabase: AT144, AT101) are less enriched and have higher Ta/La ratios (Figure

1.12). This less enriched, back-arc like signature suggests a depleted mantle source only

moderately contaminated by subduction, e.g., from earlier in the islands history. As will be

discussed in Chapter 3, based on isotopic mixing (<3% addition of sediment; Figure 3.7)

diagrams, only a small amount of sediment addition is required to achieve the observed isotopic

values. A transpressional versus compressional environment would result in lower pressure

melting that does not require a very high water content, and thus, a transpressional tectonic regime could be responsible for these large volume, tholeiitic, LREE depleted volcanics.

Komandorsky Basin spreading occurred at ~20-9 Ma, which could indicate that the spreading in this area started in the vicinity of the Chichagof Harbor and Holtz Bay volcanics on northern

Attu and moved northward toward the Komandorsky Basin with time.

The 33 Ma Casco Point (basaltic andesite pillow breccias: CP8-16B, CP98, CP916C) and

Murder Point (basalt to basaltic andesite pillow breccias and pillow lavas: MP80-2, MP10, and

MP80-8) samples (to the east of Krasni Point) differ from the 35-40 Ma volcanics (Holtz Bay and Chichagof Harbor) in that they have Sr enrichments, more enriched LREEs and HREEs, and

slightly lower εNd. The 30 Ma Sarana Bay samples (SB80-5A, SB80-6A, SB80-18, SB80-20, and

SB1A) are similar to the 33-40 Ma samples in εNd (Figure 1.14), REE concentrations (Figure

1.4), and lower Ta/La (Figure 1.12), but they have slightly enriched LREEs, larger Pb depletions,

smaller Ta, Nb, and Th depletions, and smaller Ba, U, Sr, and K enrichments (Figure 1.5). The

Casco and Murder Point samples are slightly more REE enriched than the Sarana Bay samples.

They also have depleted Ba/La, while the rest of the 33-40 Ma volcanics vary from low to arc-

like. This is most likely due to the larger volume volcanism (larger degrees of melting) and/or

less sediment/fluid addition to the north, which could indicate a more extensional environment

on the northern part of the island. Whereas the southern part of the island experienced smaller

volume (smaller degree melts) volcanism and/or more sediment/fluid addition, which may have

been due to a more compressional environment in the south. The extensional regime that may

have been beginning at this time becomes more distinctive with time.

Approximately 5 Ma after the 40 Ma pillow basalts were emplaced, the 30-35 Ma plutons

(Krasni Point and Jackass Pass on the Matthews Mountain peninsula) intruded the 40 Ma country

rock and the related 35-40 Ma volcanics erupted in Chichagof Harbor and Holtz Bay as well as

some inland locations, followed by the 33 Ma Murder Point and Casco Point volcanics and the

30 Ma Sarana Bay volcanics. Though the plutons were emplaced at approximately the same time

as the volcanics erupted, they were not exposed until later. Based on the fact that the 19 Ma

volcanics are subaerial, uplift occurred sometime before the 19 Ma volcanics were erupted. This

uplift would have resulted in the erosion that eventually exposed the plutons. However, it is not

clear how long it took for the erosion to expose the plutons. Throughout this time period,

sedimentary facies indicate that most or all of Attu remained below sea level (Gate et al., 1971).

The 30-35 Ma plutons have Ta/La ratios (Figure 1.12) that span the range from ocean ridge-like

to arc-like and do not distinctly vary systematically with location. The plutonic samples

(gabbros: AT80-61A, AT80-65; diorites: AT80-28, AT80-27) in Jackass Pass (near Matthews

Mountain) have the highest La/Sm and Sm/Yb ratios of the pluton group (0.85-0.9 and 1.21-

1.59, respectively) and overlap with the lower end of the volcanic group (0.57-1.24 and 1-1.9,

respectively) (Figure 1.11). These samples also have slightly higher on average Th/La and Ba/La

(Figures 1.12, 1.13) and smaller Ta depletions than the southern pluton samples. The plutons

have lower εNd than the volcanics (Figure 1.14), while there Sr isotopes are similar, but likely altered. The pluton’s Ba/La is higher (or altered), their Th/La (0.03-0.14) is slightly higher on

average than North Pacific MORB (0.019-0.097) and the 30-40 Ma volcanics (0.02-0.1) (Figure

1.13), their Sr isotopes are higher than MORB (Figure 1.14), and they have some of the largest

Ta, Nb, Th, Zr, and Hf depletions and Ba, U, Sr, and small Pb enrichments (Figure 1.5). Again, while some of these features might reflect alteration, not all could be due to alteration. The

location of the 35 Ma plutons farther south on Krasni Point (AT80-22, AT80-20, AT80-43A,

AT80-43, AT80-48, AT80-54, AT80-78, AT8-87, and AT80-89) (AT80-87; 29.38±2.5 and

AT80-22; 34.59±0.99 Ma) and closer to the trench may explain the stronger subduction

influence in their geochemical signature. If a more extensional regime allowed for upwelling asthenosphere to recharge the mantle wedge, then the enrichments and depletions caused by

sediment/fluid addition from subduction would be diluted by the influx of fresh mantle material.

A decrease in extension/rifting to the south plus increased proximity to sediment/fluid input from

the slab could account for the larger depletions and enrichments of the southern plutons. A

smaller degree of melting (smaller volume plutons) in the north could account for the higher

La/Sm and Sm/Yb ratios of the northern plutons.

Over all, the plutons and volcanics form a continuum in La/Sm-Sm/Yb space (Figure

1.11). The plutons have the lowest ratios, followed by the inland volcanics; then Chichagof

Harbor, Holtz Bay, and Sarana Bay; and finally, the Casco Point and Murder Point samples.

There is a fairly large range in Sm/Yb, but the La/Sm has a much smaller variation. The plutons have higher more arc-like Ba/La and Th/La (Figures 1.12, 1.13) and larger depletions and enrichments (Figure 1.5); whereas, the volcanics trend to higher LREE and HREE enrichments.

An alternative view of the formation of the 30-35 Ma plutonic rocks involves a large

amount of rotation occurring on Attu since their emplacement. Although we do not know the

exact convergence angle at this time, if the now NE trend exhibited by the normal faults, valleys,

sedimentary deposits, and the 16 Ma Albite-Granites and Late-Dikes (as mapped by Gates et al.,

1971; Yogodzinski et al., 1993), were originally formed parallel to the trench, then the plutons that are now north and south relative to the modern trench would have been approximately equidistant from the trench and aligned west and east along the trench, respectively. However,

this would require that the island had rotated ~35° counterclockwise since the emplacement of

the plutons to reach the present alignment of the island. Geist et al. (1988) calculated the rotation

of the various blocks along the Aleutian Ridge based on the rotation required to create the

observed Late Miocene to Early Pliocene summit basins. They determined that the Near Block

(Attu) had to have rotated 3° clockwise to form the Ingenstrem Depression. It should be noted that this 3° of clockwise rotation only accounts for the rotation that occurred in the last 10 Ma.

This would increase the required CCW rotation to ~38°. It is generally hypothesized that all of the blocks along the Aleutian ridge have experienced CW rotation. No one has yet proposed

CCW rotation of any of the blocks. This makes an arc front orientation that requires ~38° CCW rotation unlikely. This alignment could also not account for the differences in geochemistry seen between the two areas of plutonism. However, varying distances from the trench are known to cause differences in geochemistry due to differences in the amount of compression/extension and differences in the depth of the slab. The fact that temperature increases with increasing depth can affect the amount of fluids and melt available to produce the volcanism directly above a given part of the slab.

The Komandorsky Basin opened along a spreading axis that trends NE in the north and becomes more NNE trending in the south. This implies a spreading axis of ~45-67.5°. The ~38° alignment of the structures on Attu come close to this with Geist et al.’s (1988) proposed 3° CW rotation. However, Minyuk and Stone (2009) proposed a 50° CW rotation of Amchitka (Rat

Block) ~15 Ma, which is 30° more than that proposed by Geist et al. (1988) (20° CW). This suggests that the rotation of the Near Block could have been greater than the 3° CW rotation that

Geist et al. (1988) calculated, and it also supports the CW rotation of the blocks. If the Near

Block experienced more than 3° of rotation, this could bring the orientation of the structures on

Attu even closer to that of the Komandorsky Basin’s spreading axis. An increase in rotation of 8-

29.5° CW would bring the ~38° on Attu to the 45-67.5° of the Komandorsky Basin’s spreading axis. However, even at ~38° the extensional alignment on Attu is close enough to that in the

Komandorsky Basin to be explained by further variations with distance as are seen along the spreading axis in the Komandorsky Basin.

After a 10 Ma lull, during which the submarine sedimentary Nevidiskov, Chuniksak, and

Chirikof Formations were deposited (Gates et al., 1971), the 19 Ma Massacre Bay volcanics

were erupted (MB80-15, AT80-27A, and MR80-24). The Massacre Bay Formation also includes

the first subaerial sedimentary units (Gates et al., 1971), indicating uplift and erosion exposed the

plutons during this 10 Ma volcanic lull. These samples have low Ta/La (Figure 1.12); fairly low

REE enrichments, low Ba/La, and high εNd (Figures 1.4, 1.12, 1.14); but they have low to arc-

like Th/La (Figure 1.13), flat to slightly depleted LREEs, depletions in Ta, Nb, Th, and Pb; and enrichments in Ba, U, K, and Sr (Figure 1.5). They are very similar to the inland volcanic samples, but are less enriched than the Holtz Bay and Chichagof Harbor samples. Despite having erupted in close proximity to the earlier Sarana Bay volcanics, they are quite a bit less enriched and fall between the 30-35 Ma plutonics and the 30-40 volcanics in La/Sm-Sm/Yb space (Figure

1.11). The Massacre Bay volcanics appear to be smaller volume than the Sarana Bay volcanics, yet they are less enriched, which is inconsistent with their being produced by a smaller degree of partial melting. Instead, this could reflect a decrease in sediment/fluid addition to the mantle wedge and/or upwelling depleted mantle recharging the mantle wedge in the 10 Ma lull between the eruption of the 30 Ma Sarana Bay volcanics and the 19 Ma Massacre Bay volcanics. An increasingly more extensional environment could also facilitate this upwelling.

These are followed by the 16 Ma Albite-Granites that are very soon after intruded by the

16 Ma Late-Dikes. As suggested by Yogodzinski et al. (1993), very little time elapsed between the eruption of the Albite-Granites and their intrusion by the Late-Dikes, so we place both the dated Albite-Granites and the undated Late-Dikes at 16 Ma. These volcanics and intrusives form a NE trending belt along with sedimentary deposits (as mapped by Gates et al., 1971;

Yogodzinski et al., 1993), valleys, and normal faults, which indicates a strong extensional influence at 16 Ma.

Based on their lack of or much smaller enrichments and depletions that are typical of

subduction additions to the mantle, e.g., Ba, K, Sr, Rb, Th and Ta, and their silicic nature, the

Albite-Granites are interpreted to be the product of open source fractional crystallization and

assimilation of volcanogenic sediment or older volcanics during transit and eruption rather than

as slab involvement. The Pb isotope ratios of the Albite-Granites are distinct from the rest of the

Attu samples (Yogodzinski et al., 1993). This supports the idea that they assimilated crust during

their emplacement, particularly the increase in Pb isotope ratios with increasing silica content,

which cannot be explained by varying degrees of partial melting. Due to the very short time

between the eruption of the Albite-Granites and the Late-Dikes as well as their co-location (the

Late-Dikes intrude the Albite-Granites), we conclude that they have the same mantle source, but

different eruption/evolutionary paths, which led to their differences in chemistry.

The fact that their εNd does not change with silica content, but the Pb isotopes increase

with silica content (Figure 1.17), indicates that any assimilation would have to be of material

with εNd similar to that of the mantle source, but higher Pb isotopes. Yogodzinski et al. (1993)

interpreted this as mixing between north Pacific MORB and Pacific marine sediments. Given

that the first record of subaerial sediments coincides with the 19 Ma Massacre Bay volcanics,

207 204 Figure 1.17: SiO2 vs. Pb/ Pb for Attu. Data from Yogodzinski et al., 1993.

most of the sediments that the 16 Ma Albite-Granites erupted through would have been marine

sediments, in line with Yogodzinski et al. (1993), which indicates that the Pb isotopes of the 16

Ma Albite-Granites are most likely the result of assimilation of pelagic marine sediments (as suggested by Yogodzinski et al., 1993) that the Late-Dikes did not undergo. Alternatively, and possibly more likely, the Pb isotope signature could be due, at least in part, to the assimilation of older volcanogenic sediment that has been altered by seawater. Seawater alteration could alter a fluid mobile element like Pb without changing an immobile element like Nd. Pacific Sea water values (207Pb/204Pb = ~15.6, 206Pb/204Pb = 18.66-18.85,208Pb/204Pb = 38.60-38.91; Blanckenburg et al., 1996) are higher than those of the magmatic rocks on Attu (207Pb/204Pb = 15.40-15.517,

206Pb/204Pb = 18.20-18.901,208Pb/204Pb = 37573-38.072; Yogodzinski et al., 1993).

In summary, during this magmatic episode, we see a strongly extensional rifting chemical

signature that has been altered by fractionation and assimilation. This geochemical signature is

consistent with the extensional NE alignment of the features mentioned above, indicating that the

extensional environment played an important role in the 16 Ma magmatism.

Emplacement of the Late-Dikes was followed by another 10 Ma lull, in which the

subaerial sedimentary Faneto Formation was deposited, before the 6 Ma eruptions that formed

Matthews Mountain (AT80-30A, AT80-32, AT80-36, and AT83). The Faneto Formation is cut

by normal faults, which indicates that the tectonic environment was still extensional at this time.

Eruption of the Matthews Mountain Formation marks a significant change in chemistry, and by

inference, in the magmatic environment. Rocks of the Matthews Mountain Formation have a

distinctive calc-alkaline arc-like signature with steep LREE patterns and lower HREE than the

older samples (Figure 1.4). They have some of the most incompatible element-enriched samples

on Attu for a given SiO2 content (Figure 1.10), as was pointed out by Yogodzinski et al. (1993).

Large depletions in Th, Nb, and Ta and large enrichments in Ba, U, K, Pb, and Sr show a classic

arc-like signature (Figure 1.5). The Th/La and La/Sm values of the Matthews Mountain samples are some of the highest on Attu (Figures 1.11, 1.13). They have some of the lowest εNd on Attu

with arc-like Sr isotope ratios (Figure 1.14). They lack the obvious signs of alteration present in

the older Attu samples, so the Sr isotope ratios likely reflect magmatic values.

This change to arc-like calc-alkaline magmatism could be the result of the clockwise shift

in Pacific Plate velocity at ~6 Ma, from the cessation of Melanesian subduction due to collision

with the Ontong Java Plateau (Austermann et al., 2011). This would have resulted in a shift from

the previously more transtensional tectonic environment, which led to rifting and ocean ridge-

like signatures in previous magmatism, to a more transpressional setting that would result in a

more subduction-oriented environment and prevent the extension necessary for continued rift

related magmatism. It should also be noted that this is the only time at which there is calc-

alkaline volcanism occurring along the entire arc (Kay et al., 1990). This could also account for

the shift from tholeittic to calc-alkaline volcanism (as discussed below in the Tholeiitic to Calc-

Alkaline Evolution section).

Kiska Island

As noted above, the Kiska samples all have arc-like trace element characteristics (Figure

1.12). The older volcanics are tholeiitic, with the exceptions of the 46 Ma Murray Canyon dredge

sample (TN-182-30-003), the Sea Knob (13.57 Ma, andesite, KS-12-8), and the Gertrude Cove

rhyolite (38.54 Ma, KS-12-20) (Figure 1.6), which plot in the calc-alkaline field. These samples

are all extensively altered, so it is possible that they are actually tholeiitic, but alteration has

changed their major element chemistry sufficiently to make them appear calc-alkaline. The 5 Ma

and younger samples are calc-alkaline and resemble the 6 Ma Matthews Mountain samples from

Attu.

The Kiska volcanics are somewhat enriched in Ba, U, K, Pb, and Sr, but have lower Th,

Nb, and Ta, and lower εNd than do the rocks from Attu (compare Figures 1.5 and 1.9). These

slightly greater enrichments may reflect a more classical orthogonal subduction environment and

more sediment subduction compared to the highly oblique subduction beneath Attu. This is

supported by the lack of evidence of rifting on Kiska, in contrast to Attu. Th/La, Sm/Yb, and

La/Sm generally increase (Figures 1.11, 1.18) with time on Kiska, although in the historic

samples Th/La returns to lower values. There is also an overall decrease in εNd with time (Figure

1.18b), but it is nearly constant between 39 and 14 Ma, indicating stability in the source reservoir

after the initial arc-building event. The Sr isotope ratios (Figure 1.14) of the older samples are

generally higher than those of the younger ones, but as the young samples have some of the

87 86 lowest εNd values, we suspect that the high Sr/ Sr in the older samples results from seawater interaction.

The increase in Th/La, La/Sm, and Sm/Yb and the decrease in εNd with time suggest a

more subduction-modified source over time for both Attu and Kiska and perhaps a smaller

degree of melting, which would increase the concentration of incompatible elements. A smaller

degree of melting is supported by the observation that the volume (based on geologic map

coverage) of magmatism decreased with time as it moved north, with the majority of the island

composed of older volcanics. This more compressive environment eventually led to the tholeiitic

to calc-alkaline transition, which supports a shift to more orthogonal subduction and a more

compressive tectonic regime over time and particularly between 13-5 Ma. Unfortunately, the

lack of data points within this time frame, which is likely due a lull in magmatism rather than a

Figure 1.18: Attu and Kiska data for dated (with error bars) and age estimated samples. A) age vs. Th/La, B) age vs. εNd, C) age vs. La/Sm [normalized to C1 chondrite (ratio divided by 1.62) (McDonough and Sun, 1995)], and D) age vs. Ta/La. Average values of MORB depicted by black lines and error ranges by blue fields as determined by White and Klein (2014). (Includes Attu data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984. Includes Kiska data from George et al., 2003.)

lack of samples, makes it hard to narrow down the time frame. However, the shift in plate motions at ~6 Ma (Austermann et al., 2011) favors the lower end of the age range.

A smaller extent of melting would not explain the change in εNd with time, which

requires a change in the source, most likely due to slab contribution. Although the Ba/La of the

Kiska samples varies considerable, there appears to be no systematic variation with time (the ratios in the altered older rocks are suspect), indicating that the hydrous fluid contribution remained roughly constant. This suggests that the change resulted from a greater incorporation of the slab-melt component into the mantle wedge.

Tholeiitic to Calc-Alkaline Evolution

Miyashiro (1974) proposed that water from the slab is a critical factor in determining whether magmas evolve along the tholeiitic or calc-alkaline trends because water is a strong controlling factor of the oxygen fugacity of the mantle wedge. As the water dissociates into H2

and O2, the oxygen fugacity increases, favoring early crystallization of oxides and calc-alkaline evolution. However, more recent studies (Frost and Ballhaus, 1998; Evans, 2012) suggest that the water itself is not the oxidant, but instead, it carries oxidants such as oxidized sulfur and iron from the slab into the mantle wedge. As time progresses, the mantle wedge can become depleted in basaltic components and become less capable of producing magmas (Miyashiro, 1974).

Eventually, larger amounts of water are necessary to create new melts; and the addition of more water results in higher oxygen fugacities. Higher oxygen fugacities lead to early crystallization of Fe oxides, while higher water contents suppress plagioclase precipitation; and together this produces calc-alkaline rather than tholeiitic evolution. Though depletion of the mantle lowers the concentration of REEs, the smaller percent melts result in enrichment of REEs in calc-alkaline

rocks. High pressure also suppresses plagioclase precipitation. Thick low-density arc crust will favor crystallization at greater depth, and thus, as arc crust thickens with time, calc-alkaline

rather than tholeiitic evolution will be favored. In this respect, it should also be noted that the 6

Ma Matthews Mountain and Massacre Bay volcanics have reasonably high Ba/La, despite being

young and unaltered, indicating the presence of a fair amount of water. Subsequent work by a

number of investigators has largely substantiated Miyashiro’s (1974) speculation and provided

additional insights into the distinction between tholeiitic and calc-alkaline evolution.

Sisson and Grove (1993) approached the cause of the calc-alkaline and tholeiitic trends

experimentally. They showed that hydrous basalts have higher oxygen fugacities and crystallize

the iron oxide phase early, resulting in Fe-depletion and silica and alkali enrichment, i.e., the

calc-alkaline trend. Anhydrous basalts have lower oxygen fugacities and do not crystallize iron

oxides, resulting in Fe-enrichment and only modest increases in silica, i.e., the tholeiitic trend.

The presence of water also destabilizes early plagioclase, but does not destabilize oxides as

much, which explains why plagioclase is an early phase in tholeiites and not in calc-alkaline

rocks and why oxides precipitate early in calc-alkaline rocks. Sisson and Grove’s (1993)

experiments with high-alumina Aleutian basalts indicate that at 2 kb, with an NNO buffered

mantle, and 4.5 wt% H2O, Fe-oxide is at or near the liquidus, i.e., calc-alkaline. At 4-6 wt% H2O

and typical subduction zone oxygen fugacities, calc-alkaline melts can be created.

Zimmer et al. (2010) further investigated the contribution of water to the evolution of the

calc-alkaline trend. They developed a quantitative index of Fe enrichment, the Tholeiitic Index

(THI), to distinguish between tholeiitic and calc-alkaline evolution. Their tholeiitic index looks

at tholeiitic and calc-alkaline solely in terms of Fe-enrichment and Fe-depletion, respectively.

They define THI = Fe4.0/Fe8.0, where Fe4.0 is the average Fe* of samples with MgO = 4±1 wt%

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and Fe8.0 is the average FeO* of samples with MgO = 8±1 wt%. Based on this equation, they

found that magmas with THI>1 become enriched in iron during differentiation from basalts to

andesites and are tholeiitic, while magmas with THI<1 are calc-alkaline. Furthermore, they point

out that the dividing line on the Miyashiro (1974) diagram is at THI=0.9 instead of THI=1,

suggesting the need to update Miyashiro’s diagram. They found that for Aleutian lavas as well as

global arc samples, THI and H2O are negatively correlated, i.e., THI increases with decreasing water content, with THI=1 corresponding to a water content of 2 wt%. To further determine the amount of control the water content of the magma has on the calc-alkaline trend, they investigated other factors that have been suggested to result in the calc-alkaline trend. They concluded that suppression of plagioclase and crystallization of Fe-rich minerals early on are primary factors in controlling the trend; whereas, amphibole tends to crystallize too late in the sequence to have much effect on the Fe-depletion, and early amphibole does not have a high enough Fe content to cause sufficient Fe-depletion to result in the calc-alkaline trend.

The role of pressure on the crystallization sequence has been suggested as a major factor in controlling the calc-alkaline trend. At high pressures, even low water contents suppress plagioclase crystallization, which results in a larger proportion of Fe-bearing mafic minerals crystallizing, buffering Fe content. However, at low pressures, insufficient water can be dissolved to suppress early plagioclase crystallization, resulting in a tholeiitic trend. Zimmer at al. (2010) determined that the effect of pressure on the THI was ~20% increase when crystallization occurs at 1 GPa versus 0.2 GPa, which is enough to cause scatter around the THI versus H2O array, but not enough to alter the array. From this investigation, they conclude that

water content is the primary factor causing the calc-alkaline trend, not pressure. The effects of

fO2 are less clear as fO2 increases with water content, but Zimmer et al. (2010) were still able to

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determine that the THI is related to variations in water content at elevated fO2 (>FMQ) and that

water content was the primary factor and fO2 a secondary factor. The last factor they investigated

was the effect of major element composition, e.g., varying FeO content or CaO/Al2O3, on THI.

They determined that it had a 5-10% effect on the THI and was also a secondary factor. In

summary, Zimmer et al. (2010) found that water was the primary factor in determining the calc-

alkaline versus tholeiitic trend, with tholeiites having H2O <2 wt% and THI>1. Factors such as

pressure, fO2, and major element composition had only a secondary effect. Unfortunately, using

the THI of Zimmer et al. (2010) requires multiple samples from the same magma series and an

MgO range of 4-8 wt%. In the case of Kiska and Rat and some of the formations on Attu, we do

not have sufficient samples to perform the calculation. In the case of Attu, even when there are a

sufficient number of samples from a single formation, we are missing one or both ends of the

MgO range required for the calculation. Therefore, in this study, we have used the calc-alkaline

versus tholeiitic classification of Miyashiro (1974).

Kelley and Cottrell (2009) found that Fe3+/ΣFe increases from ridges to back arcs to arcs

and correlates linearly with water content as well as the fluid mobile elements in slab fluids. The oxygen fugacity of the subducting plate is higher than that of the mantle and can supply S6+ to

the mantle wedge, most likely through slab fluids, which results in higher oxygen fugacities in

the mantle wedge. Each S6+ can oxidize 8 Fe2+ to Fe3+. This mechanism avoids the need to

dissociate the water and remove the resultant H2 to increase fO2.

The study of Brounce et al. (2015), which focused on the Mariana Arc, builds on the investigation of mantle wedge oxidation. They found that within 1-4 Ma of arc initiation, arc volcanics reach mature arc (more evolved) values of concentrations of fluid mobile elements and oxygen fugacity. As these two factors correlate, they concluded that slab fluids play a dominant

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role in determining the oxygen fugacity of the mantle wedge. This transition in chemistry likely

signals the transition from initial decompression melting to fluid induced melting. As the

subduction zone becomes well developed, the increase in oxygen fugacity will cease as recharge

from the deep mantle balances the fluid input from the slab. Models suggest that a 14 times

increase in oxygen fugacity can be achieved in as little as 2-4 Ma.

Finally, Plank and Langmuir (1988) noted that the major element chemistry of arc

magmas correlated with the thickness of the arc crust/lithosphere. In particular, Na2O increases

and CaO decreases with crustal thickness when corrected for fractional crystallization. They

argued that this was a consequence of crustal thickness determining the height of the melting column, with thick “crust” resulting in shorter melting columns and smaller extents of melting and, consequently, higher Na2O. As calc-alkaline magmas are characterized by higher Na2O/CaO

than tholeiitic magmas, thick crust and small extents of melting will favor calc-alkaline

volcanism.

Kay et al. (1982) proposed a tectonic control on tholeiitic and calc-alkaline magmatism.

According to their model tholeiitic centers are larger and are mainly basaltic due to their placement at the edge of or between tectonic blocks where stresses are less compressional, allowing for easier passage of the magma to the surface and shallow closed system differentiation. Calc-alkaline centers are smaller, more andesitic, and tend to occur in the middle of tectonic blocks. This placement produces a more compressional environment that prevents easy transport of magma to the surface and results in deeper differentiation and a larger percentage of intrusives than with tholeiitic centers. The greater pressure reduces the stability of plagioclase and enhances that of amphibole, hence limiting iron enrichment during fractional

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crystallization, resulting in the magma evolving to a calc-alkaline rather than a tholeiitic

composition. More fractionation at depth also increases the water content.

On both Attu and Kiska, as well as in other parts of the arc (Kay et al., 1982), the older

samples are predominantly tholeiitic, while the youngest samples are calc-alkaline. This model

(Kay et al., 1982), if applicable to the Western Aleutians, which suggests a change from an

extensional environment to a compressional one with time. On Attu, we see evidence of

extensional tectonics in the older rocks until as recently as 16 Ma; however, there is no evidence

of this extension in the 6 Ma Matthews Mountain samples. In fact, we see a distinct arc-like

signature at 6 Ma. Thus, a shift to a more compressive tectonic environment may have caused or

contributed to the shift from tholeiitic to calcalkaline evolution, consistent with the model of

(Kay et al., 1982).

On Attu, we see evidence of at least an intermittently extensional regime in the older

rocks until as recently as 16 Ma; however, there is no evidence of this extension in the 6 Ma

Matthews Mountain samples. In fact, we see a more distinct arc-like signature at 6 Ma. The 6 Ma shift in plate motion (Austermann et al., 2011) resulted in a more perpendicular subduction at

Attu and a more classic subduction zone setting, which may also have resulted in a more compressive environment, forcing the shift to higher pressure and smaller degrees of melting,

resulting in calc-alkaline magmatism.

The question thus becomes which factors are affecting the tholeiitic to calc-alkaline

transition on Attu? Most likely, all of these factors: a more water-rich oxidized mantle source,

greater crustal thickness, and a shift to a more compressive tectonic regime contributed to the

shift from tholeiitic to calc-alkaline magmatism on Attu.

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The Attu samples do not show a chemical trend with age (Figure 1.18), but rather a

change between 16 and 6 Ma to a new chemical regime. In the case of the Kiska samples, a

change in Th/La occurs at 13 Ma, while the change in La/Sm was more gradual. This could

indicate a slower change in plate motion geometry, rather than an abrupt shift, which would be

more realistic, and is consistent with a gap in the magmatic rock record on Attu between 16 Ma

and 6 Ma. The same change in tectonics could be responsible for the shift in geochemistry on

both Kiska and Attu during this time period.

The other reason may be a lag in the timing of the tectonic transition due to the more oblique nature farther to the west. Since this area started out more oblique, a larger change would be required before it reached the point where the change was distinct. In summary, the youngest

samples on Attu and Kiska are the most enriched and show the most subduction influence. This

change in chemistry is most likely due to the change in environment discussed above, which was

caused by the change in Pacific Plate motion at 6 Ma (Austermann et al., 2011); however, a

change in motion several million years earlier cannot be ruled out without more data points

between 16-6 Ma. It appears that the change in plate motion had a greater effect on the western

Aleutian Arc, likely due to the oblique nature of subduction making the area more sensitive to

changes in subduction angle.

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Figure 1.19: Tectonic evolution model of the North Pacific-Bering Sea region. A) Initial Aleutian Ridge building event at 46 Ma (Kiska) and 40 Ma (Attu). Bowers Ridge volcanism starts at approximately this time. Compressive Shirshov Ridge already present. Attu and Kiska not yet above sea level. B) 41-30 Ma Chichagof Harbor, Holtz Bay, Plutons. 30 Ma Serana Bay, Murder Point, Casco Point. Gates et al. (1971) sedimentary submarine basement series. Bowers Ridge volcanism Continues. Kiska volcanism from 39-31 Ma. C) 19 Ma Subaerial Massacre Bay. 30-19 Ma Gates et al. (1971) sedimentary submarine Nevidiskov, Chuniksak, and Chirikof Formations were deposited. Bowers Volcanism ends 26-22 Ma. Komandorsky spreading begins 20 Ma. Kiska volcanic lull from 31-15 Ma. D) 16 Ma Subaerial Albite-Granites and Late-Dikes. Komandorsky spreading ends at 9 Ma. Kiska volcanism 15-13 Ma. E) 6 Ma Subaerial Mathews Mountain. Kiska volcanism 5-0 Ma.

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107

Figure 1.19 (Continued):

108

Figure 1.19 (Continued):

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are depicted as different colored fields. (41.5±2 Ma amphibolite from Vallier et al., from Ma amphibolite 1983) fields. (41.5±2 colored as different are depicted Figure 1.20: Map depicting the magmatic evolution of Attu Island. Different episodes of volcanism volcanism episodes of Different Attu Island. evolution of magmatic the Map 1.20:depicting Figure

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TECTONIC MODEL FOR ATTU AND KISKA ISLANDS

Attu Island

Our model (as depicted in Figures 1.19 and 1.20), based on a wider range of Attu samples, better age constraints, and the progression of tectonic events as modeled by Scholl

(2007), differs somewhat from that of Yogodzinski et al. (1993). Based on additional analyses

and the earlier amphibole K-Ar age of 41.5±2 Ma on an amphibolite (Vallier et al., 1983;

Pickthorn and Vallier, 1991), we propose that the history of Attu begins well before 40 Ma with

the voluminous eruption of arc-like tholeiites such as sample AT80-76 at ~ 40 Ma (Figures 1.19

and 1.20). These tholeiites are slightly LREE enriched (Figures 1.4, 1.11), have very low εNd (the lowest observed on Attu), low Ta/La (0.01; Figure 1.12), and alkali and alkaline earth enrichments (although the extent to which the latter are due to secondary alteration remains unclear), but lack the Pb enrichment observed elsewhere in the Aleutians (Figure 1.5). The lack of Pb enrichment may have been the result of subducting relatively continental sediment-poor

material initially or a regional characteristic of the mantle in this region. This proposed voluminous arc-related volcanism is consistent with Vallier et al.’s (1994) proposal of early voluminous arc-like magmatism due to more orthogonal subduction. The established subduction signature of the 40 Ma arc-like tholeiitic magmas and the volume required to build the ridge to this point (Jicha et al., 2006) indicates that subduction in the western Aleutians was well established by 40 Ma, which indicates that 50 Ma (consistent with Scholl, 2007) or even earlier

~55-50 Ma (consistent with earlier dates on Adak and Kiska) would not be an unreasonable estimate for the timing of the initiation of subduction in the western Aleutians and likely the entire arc.

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Next, the tholeiitic 35-41 Ma volcanics (Chichagof Harbor and Holtz Bay) and the 30-35

Ma plutons were emplaced (Figures 1.19 and 1.20). The volcanics are more REE-enriched overall than the intrusives rocks. This could be due to differences in the amount of partial melting. The plutons are more enriched in K, Sr, and Ba, more depleted in Nb, Ta, and Th, and more enriched in εNd, i.e., they have a stronger subduction signature than the volcanics. As the plutonics and volcanics have approximately the same range of silica contents, the stronger subduction signature in the south is most likely due to the proximity to the subducting slab. The melt being erupted to the north interacted with a deeper already fluid/melt depleted slab, and thus, it received less influence from the slab.

The 33-30 Ma volcanics that followed are more variable (Figures 1.19 and 1.20). The 30

Ma Sarana Bay samples (SB80-5A, SB80-6A, SB80-18, SB80-20, and SB1A) are similar to the

Chichagof Harbor and Holtz Bay REE enrichments (Figures 1.5, 1.12), which is consistent with their location to the south of the Chichagof Harbor samples. However, the 30 Ma Sarana Bay samples have smaller Th, U, Nb, and Ta depletions and Sr, K, and Ba enrichments than the 35-40

Ma Chichagof Harbor and Holtz Bay samples. The Sarana Bay samples are likely a continuation of the Chichagof Harbor volcanics (CH samples) that have become slightly less subduction- related with time, indicating that the source has been diluted by an influx of more depleted mantle, diluting the subduction signature, i.e., a backarc signature. The 33 Ma Casco Point (CP8-

16B, CP98, and CP916C) and Murder Point (MP80-2, MP10, and MP80-8) volcanic rocks

(located just east of the Krasni Point plutons and country rock) are more LREE-enriched, which is most likely due to a smaller degree of partial melting to the south. They also have lower (more enriched) εNd. They also have lower Th, U, Nb, and Ta concentrations and higher Sr, K, and Ba concentrations than the Sarana Bay samples (Figures 1.4, 1.5). These differences are likely also

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due to the influx of depleted mantle to the north, which diluted the arc signature for the Sarana

Bay volcanics. Thus, the fact that the Casco and Murder Point volcanics and the Sarana Bay

volcanics erupted at approximately the same time, but with different geochemical signatures, can

be explained by a combination of more fluid/melt input from the shallower part of the slab closer

to the trench, with decreasing input as the slab travels deeper and is dehydrated (melts that erupt

to the north), and incipient intra-arc rifting occurring to the north (Sarana Bay), with larger

volume melts caused by decompression melting reducing/diluting the subduction signature while

arc volcanism occurs to the south (Casco and Murder Points). Thus, by at least 33-30 Ma, frontal

arc volcanism appears to have been active in the south, while rifting was occurring behind the

main arc front (Figures 1.19 and 1.20). This combination of an active main arc with back-arc

rifting has been seen in other arcs around the Pacific Rim, such as the Izu-Bonin-Mariana Arc,

which has much more advanced back arc rifting (Arculus et al., 2015; Reagan et al., 2010) and

the Tonga-Kermadec Arc (Ewart et al., 1977).

Scholl (2007) argued that Bowers Ridge likely initiated subduction simultaneously to or

immediately after subduction along the Aleutian Arc. The ages reported by Wanke et al. (2012)

indicate that Bowers Ridge ceased to be active around 26 Ma and Bowers seamounts around 22

Ma. However, the dated samples are from the crest of the ridge and the relief of the ridge, so they

do not provide good constraints on the initiation of activity. Allowing for several million years to build to the height from which these samples were taken, this would put the initiation of Bowers

Ridge at roughly 40 Ma, albeit with considerable uncertainty (Figures 1.19 and 1.20). Bowers

Basin spreading preceded that in the Komandorsky Basin, which based on heat flow, probably began at approximately the same time as activity on Bowers Ridge ceased (Cooper et al., 1992)

(Figures 1.19 and 1.20). This places the lifetime of Bowers Ridge and Bowers Basin activity

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from ~ 40 Ma to 22 Ma, overlapping with the initial history of Attu and ceasing just before the

opening of the Komandorsky Basin (20-9 Ma) (Figures 1.19 and 1.20) (Rubenstone, 1984;

Valyashko et al., 1993; Scholl, 2007). The cessation of compression in the Shirshov Ridge area is thought to have occurred at the latest by the time spreading in the Komandorsky Basin started

(Scholl, 2007; Sukhov et al., 2011), approximately the same time Bowers Ridge became inactive

(Figures 1.19 and 1.20).

After 30 Ma, there was a 10 Ma lull in volcanism before the 19 Ma Massacre Bay volcanics erupted (Figures 1.19 and 1.20). A similar lull occurred on Kiska. This period corresponds to the youngest ages for Bowers Ridge, suggesting that the compressive plate motion in the area was taken up by Bowers Ridge subduction at this time.

The older subaerial Massacre Bay volcanic rocks erupted around 19 Ma, ending the lull

(Figures 1.19 and 1.20). The La/Sm (0.87-1.24) and Sm/Yb (1.4-1.9) (Figure 1.11) of the

Massacre Bay volcanic rocks fall between the 30-35 Ma plutons and the 40-30 Ma volcanics,

overlapping with the inland volcanics. Thus, the Massacre Bay volcanics represent the depleted

end member of the range of compositions of Attu’s volcanics. This is likely due to the fact that

the Massacre Bay volcanics are smaller volume than those to the north (CH, HB, and SB), and

although they are larger volume than those to the south (CP and MP), the mantle to the north has

been recharged by upwelling depleted mantle due to extension. It is notable that although the older Sarana Bay volcanics are only slightly farther north, they have less prominent enrichments and depletions. The depletions and enrichments of the Massacre Bay volcanics are similar to those of the older Chichagof Harbor and Holtz Bay volcanics, although the Ta depletions of the

Massacre Bay samples are large and arc-like. The combination of Massacre Bay volcanics having the lowest La/Sm and Sm/Yb, but more prominent enrichments and depletions is likely

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the result of the upwelling depleted mantle to the north, caused by increased extension, being re-

enriched by an increase in the sediment/fluid input with time.

The advent of the small volume 16 Ma magmatism on Attu (Figures 1.19 and 1.20)

marks a change in regional tectonics and geochemical characteristics to a more extensional environment. The εNd (Figure 1.14) of the 16 Ma magmatism does not differ from that of other

time periods on Attu, eliminating the possibility of a change in source. The 16 Ma Albite-

Granites and Late-Dikes have smaller relative Ba, U, and K enrichments and smaller Th and Ta

depletions (Figure 1.5), but they are more incompatible element-enriched overall than previous

magmatism. The Albite-Granites are more enriched overall than the Late-Dikes, which is likely

the result of fractional crystallization. As Yogodzinski et al. (1993) suggested, it is possible that

the mafic intrusives are the crystal cumulates of the basaltic magmas that resulted in the Albite-

Granites. Notably, the Albite-Granites define an inverse correlation between La/Sm (1.36-2.84)

and Sm/Yb (1.15-1.96) (Figure 1.11). La/Sm increases and Sm/Yb decreases with SiO2 (Figure

1.10) among these samples, which is difficult to explain by anything other than amphibole

fractionation. In addition, the higher Pb isotope ratios (Yogodzinski et al., 1993) of the Albite-

Granites and the increasing Pb isotope ratios with increasing silica suggests this process was accompanied by assimilation of marine sediments. As discussed in the previous section, this is most likely due to the assimilation of older seawater altered volcanogenic sediments.

The alignment of the 16 Ma magmatism along a NE-SW trend that is parallel to the

sedimentary belt, indicates that the 16 Ma units were emplaced during a rifting event (Figures

1.19 and 1.20). Intrusion of magma into well-cooled crust following a long hiatus in magmatism

in this area may explain the extensive fractional crystallization. This rifting may be unrelated to

Aleutian subduction, and instead, may be related to back arc spreading and the opening of the

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Komandorsky Basin along a NE-SW axis (Scholl, 2007; Valyasho et al., 1993) during a particularly extensional episode in the Western Aleutians. Scholl (2007) proposed that the initiation of spreading in the Komandorsky basin is linked to the tearing of the Pacific slab west of Buldir (Creager and Boyd, 1991), or beneath Attu, at approximately the time of Komandorsky

Basin Spreading (Levin et al., 2005), which allowed for the upwelling of asthenosphere

(Cormier, 1975) in the area of the Komandorsky Basin and could explain the unusually high heat flow of the basin.

Sometime after the 16 Ma magmatism took place, rotation of the Near Island Block occurred. On Amchitka, a 50° clockwise rotation is proposed at ~15 Ma (Minyuk and Stone,

2009). The amount by which Attu rotated is discussed in detail in the previous section. Attu could have rotated by 3-50° clockwise (Minyuk and Stone, 2009; Geist et al., 1988) at approximately the same time as the rotation of the Rat Block occurred, ~15 Ma.

After another 10 Ma lull in magmatism, the Matthews Mountain and adjacent Massacre

Bay Formation volcanics erupted around 6 Ma (Figures 1.19 and 1.20). This magmatism is small volume, calc-alkaline (Figure 1.6), and arc-like, with low εNd (Figure 1.14), higher Th/La (0.12-

0.14) and Ba/La (33.77-60.15) (Figure 1.13); enrichments in Ba, U, K, Pb, and Sr; depletions in

Th, Nb, and Ta (Figure 1.5); the most enriched LREE besides the Albite-Granites, and steep

LREE with lower HREE than other Attu samples (Figure 1.4). This most recent magmatic event is the most typically calc-alkaline and arc-like on Attu, and its small volume arc-like calc- alkaline nature is likely due to a shift to a transpressional regime caused by the clockwise change in Pacific Plate motion at ~6 Ma from the clogging of the Melanesian subduction zone by the

Ontong Java Plateau (Austermann et al., 2011) or from the collision of westward moving

Aleutian terranes with the Kamchatka trench, though somewhat later than modeled by

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Yogodzinski et al. (1993). This more compressive environment combined with the thicker crust now in place on Attu, due to the addition of large volumes of previous magmatism, could result in magmas stagnating at greater depths (Miyashiro, 1974), inhibiting magma ascent, and resulting in deeper differentiation where plagioclase crystallization is suppressed, and consequently, calc-alkaline characteristics are developed. In addition, Matthews Mountain has relatively high Ba/La (Figure 1.13). As Ba is fluid-mobile and La is not, this suggests a higher amount of water input to the mantle wedge, which, as discussed earlier, strongly favors calc- alkaline evolution.

The pattern of compressive subduction signatures to the south combined with extension and mantle upwelling to the north throughout the history of Attu, which is absent on Kiska, indicates a changing stress pattern through time. Initially, we have voluminous subduction volcanism around 40 Ma that forms a NE trending band of volcanism (e.g., the 41.5±2 Ma amphibolite dated by Vallier et al., 1983 and the sheeted dike complex in Chichagof Harbor), indicating that rifting had already started at this time. Then, the Chichagof Harbor and Holtz Bay volcanics (41-35 Ma) and some plutons that exhibit a much weaker subduction influence erupted, which together with their alignment, suggests extensional intra-arc rifting, causing the upwelling of depleted mantle. At 30 Ma, the Murder Point and Casco Point volcanics exhibit a stronger arc signature to the south, while the Sarana Bay volcanics to the north are more depleted and have smaller depletions and enrichments, suggesting rifting on northern Attu. This depicts a complex stress structure for Attu at 30 Ma with spreading occurring behind the subduction front, i.e., the initial stages of a back-arc basin. However, after 30 Ma, there is a 10 Ma lull in all volcanism, and when activity returns at 19 Ma, the Massacre Bay volcanics, though they have lower La/Sm and Sm/La (Figure 1.10), again have strong arc-like element depletions and

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enrichments, indicating that the upwelled mantle is being re-enriched with subduction

fluids/sediments/melts. Around 16 Ma, we see distinct evidence of extension and rifting in both

the Albite-Granites and the Late-Dikes that cross-cut them. This coincides with the opening of the Komandorsky Basin and the extensional environment required by this opening. This was followed by another 10 Ma hiatus in volcanism, after which we have only strongly arc-like calc- alkaline magmatism (6 Ma). Thus, Attu experienced episodes of intra-arc spreading, but this never progressed far enough to create a back-arc basin.

Kiska Island

The tectonic picture for Kiska is simpler than that of the more western Attu (Figure 1.19).

Chemically, Kiska is intermediate in geochemical characteristics between Attu and Adak, which is consistent with its intermediate geographical placement and supports the conclusion of

Yogodzinski et al. (1993) that Attu has always been in its present relative position along the arc,

rather than having traveled westward along the trench (Rostovtseva and Shapiro, 1998). Kiska maintains a strong subduction zone signature through time. The 46 Ma Murray Canyon dredge sample has experienced seawater alteration. This combined with its LREE depleted nature indicates that the sample is likely tholeiitic. Taking this into consideration, based on its low

Th/La, high εNd, and lack of Pb enrichment, the initial voluminous ridge-building event in the

Kiska area was tholeiitic with a relatively small sediment contribution. This may indicate a more

orthogonal and/or compressive subduction environment. The age of this sample supports the case

that arc volcanism in the Western Aleutians began at least several million years before the oldest

samples exposed on Attu (Figure 1.19). This may be supported by the potentially ~50 Ma age of

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a sample on Adak (Rubenstone, 1984). However, if volcanism began earlier, it implies lower magma production rates over the long term or substantial erosive losses.

After the initial outpouring, Kiska began to build up as a localized island (Figure 1.19).

Magmatism in this period was tholeiitic with enrichments and depletions characteristic of subduction-related magmatism and fluid and slab influence signatures, but with the same range of slightly depleted to slightly enriched REEs as seen on Attu and intermediate between MORB and Adak. This remains the case until 5 Ma (Figure 1.19) when magmatism becomes calc- alkaline and exhibits the steeper LREE patterns of the eastern Aleutians. This could be the result of the previously described Pacific Plate motion change at ~6 Ma (Austermann et al., 2011) that caused the change to calc-alkaline volcanism on Attu. Once again, this more compressive and more perpendicular subduction resulted in smaller volume, calc-alkaline magmatism similar to that seen in the central Aleutians. The occurrence of this shift and the resumption of magmatism at nearly the same time, after a brief hiatus, on both islands despite the distance between them indicates that the change in plate motion had a large impact on at least the western Aleutians and was likely significant enough to be seen in other places around the Pacific Plate as well.

On a larger scale, the hiatuses in magmatism seen on Attu and Kiska roughly coincide in time, indicating a condition that possibly affects the western Aleutians in general. If the tectonic processes in the vicinity are considered, we can develop a wider picture of the causes of these hiatuses (Figure 1.19). The Komandorsky Basin opened from ~20-9 Ma, and Bowers subduction ended by ~22 Ma. The earliest large hiatus on Attu (30-20 Ma) coincides with the youngest, and only dated, volcanic episode on Bowers Ridge (Figure 1.19). This may indicate that shifts in stress in the Bering Sea area caused a shift in subduction from the western Aleutians to Bowers

Ridge. The next large hiatus (16-6 Ma) coincides with the opening of the Komandorsky Basin

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(Figure 1.19), indicating that the stress in the vicinity of the western Aleutians at this time was

largely extensional. This is also seen in the rifting signature of the 16 Ma volcanics that preceded

the hiatus. On Kiska, this hiatus occurred from 13-5 Ma, and there is no evidence of rifting

(Figure 1.19). This could indicate that the influence of the extensional event reached the Kiska region slightly later and was not strong enough to cause rifting, but was strong enough to cause a hiatus in subduction induced volcanism. As the extension shifted farther north, volcanism on

Attu ceased, and the formation of the Komandorsky Basin accommodated the extension (Figure

1.19). When the stresses changed again, Komandorsky Basin spreading ceased and arc volcanism resumed on Attu (6 Ma) (Figure 1.19). These same hiatuses are seen on Kiska, but the volcanism on Kiska does not show the extensional regime (rifting volcanics), only the lack of compression (hiatuses) (Figure 1.19). This indicates that the extension was more severe to the west, and Kiska never started to form a back arc basin. These same hiatuses are seen in the central and eastern parts of the arc as well, e.g., Adak and the Delarofs, (Schaen et al., 2016;

Scholl et al., 1976; Kay et al., 1982, 1994; Citron et al., 1980; Fournelle et al., 1994; Jicha et al.,

2006). This indicates that the changing tectonics in this part of the Bering Sea had effects far

reaching enough to affect subduction along the entire arc. However, except for the hiatus in

volcanism, these effects are less severe farther east, indicating that the effects were perhaps less

intense with distance from the main tectonic event.

Over the history of Kiska, volcanism has slowly migrated northward ~35 km to the

present site of Kiska Volcano on the northern most tip of the island. This shift in volcanism

could be related to subduction erosion wearing away the overriding plate, and thus, the trench

migrating northward. If the upper plate is being eroded and the pieces subducted, we should see a

shift in the chemistry of the later volcanism as this new component is added to the subduction

120

system. The addition of this oceanic crustal material should decrease the Sr isotopes and increase

the εNd of the newer volcanics. We do see a decrease in the εNd of the younger volcanics; however, the 87Sr86Sr is actually lower. This decrease in Sr could simply be a reflection of the

lack of seawater alteration in the younger less altered samples; whereas, the older samples are

suspected to have been altered by seawater and have Sr isotopes that have been shifted to higher

values. If it is the base of the crust that is being eroded, it could even be eroding material that has not been altered by seawater. This makes our largest evidence for subduction erosion the shift in the volcanic line as we may not be able to distinguish the process geochemically.

SUMMARY

Through the addition of new geochemical and geochronologic data, this study has built on previous studies of Attu Island to gain a more detailed understanding of its initiation and evolution through time. In addition to the study of Attu, new geochemical and geochronologic data has been used to revise and expand what little work had previously been done on Kiska

Island. This allowed for the comparison of the two islands and an investigation into how their placement along the Aleutian Arc affected their initiation and evolution. The following

conclusions have been reached in this study.

1) On Attu, the oldest (40 Ma) large volume tholeiitic subduction volcanism exhibits

enrichments in K, Ba, and Sr (possibly altered), and depletions in Nb and Ta, resulting

from the fact that early subduction had a large initial flux of water from the subduction

initiation event. The 40-35 Ma Chichagof Harbor and Holtz Bay volcanics are also large

volume tholeiitic magmas, which are more enriched than the inland samples and the 40

Ma country rock samples despite their large volume. The 33 Ma Casco and Murder Point

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volcanics to the south are more enriched in LREE and HREE and in Ba, U, Sr, and K,

which could be due to lower extents of melting in the mantle. The 30 Ma Sarana Bay

volcanics, located just south of the Chichagof Harbor and Holtz Bay volcanics, are

smaller volume than the Chichagof Harbor and Holtz Bay volcanics, but are slightly less

enriched. This is likely caused by the upwelling of depleted mantle, which would dilute

the subduction enriched mantle wedge. This upwelling is the first clear sign of a more

extensional environment on Attu, though it could have started as early as 35 Ma, and

indicates the beginning of incipient intra-arc rifting on the northern part of the island, i.e.,

back arc basin formation. The 35-30 Ma plutons are the least LREE and HREE enriched

overall on Attu (even more depleted than their 41 Ma host rock) and form a depleted

endmember. The northern plutons (Jackass Pass) are more LREE and HREE enriched

than the southern plutons (Krasni Point). However, it should be noted that the volumes of

the plutons can only be judged by the exposed portions, and thus, the relative sizes of the

plutons are uncertain. The fact that the southern plutons exhibit stronger enrichments and

depletions is likely due to their proximity to the trench leading to more enrichment from

subduction fluids/melts that were removed from the slab at shallower depths. The

northern plutons are above the deeper, already dehydrated slab, and they also may have

experienced dilution from upwelling depleted mantle due to the more extensional

environment slightly farther north. Over all, the plutons and volcanics form a continuum

in La/Sm-Sm/Yb space (Fig. 1.10). The plutons are the most depleted, followed by the inland volcanics; then Chichagof Harbor, Holtz Bay, and Sarana Bay; and finally, the

Casco Point and Murder Point samples.

122

At 19 Ma, volcanism becomes subaerial for the first time. They are smaller

volume than the Sarana Bay volcanics, but are less enriched, indicating a continuing

extensional environment resulting in upwelling depleted mantle. There could also be a

decrease in the amount of fluid/sediment being subducted at this point in arc history. The

geochemistry of the 16 Ma Albite-Granites and Late-Dikes is consistent with the

extensional NE alignment of sedimentary, topographic, and tectonic features in the area,

indicating that the extensional environment played an important role in the 16 Ma

magmatism. Their enriched nature is caused by a combination of fractional crystallization

and the assimilation of older marine sediments. The 6 Ma Matthews Mountain volcanics

are calc-alkaline and have a strongly arc-like signature, similar to other Aleutian samples.

The calc-alkaline nature of these younger volcanics is likely due to a combination of a

more compressive environment and a thicker crust causing deeper melting and more

water influx altering the oxygen fugacity of the mantle wedge.

2) Bowers Ridge volcanism ceased at 26-22 Ma at approximately the same time the

magmatic lull on the Aleutian Ridge ended to the west. The incomplete intra-arc rifting

event that resulted in the Albite-Granites and Late-Dikes as well as the ocean ridge-like

signature in the 33-35 Ma volcanics on the northern part of Attu is possibly the southern

extension of the rifting that formed the Komandorsky Basin from 20-9 Ma. As the

extension on Attu started before 20 Ma, as evidenced by the geochemistry of the 33-35

Ma volcanics, this indicates that the extension migrated northward toward the

Komandorsky Basin with time. The end of spreading in the Komandorsky Basin roughly

coincides with the end of the lull in volcanism on the Aleutian Ridge and a shift in plate

motion, indicating a transition from a more extensional to a more compressional regime

123

on the Aleutian Ridge at this time. The change in the nature of volcanism at 6 Ma is

likely the result of the change in Pacific Plate motion that occurred at ~6 Ma.

3) On Kiska, the geochemistry is arc-like and falls between that of the volcanic rocks on

Attu and Adak in composition. Over time, the magmatism on Kiska displays a shift to

lower degrees of melting, and a more compressive environment that eventually led to the

calc-alkaline historic volcanism.

124

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White, W. M., and E. M. Klein (2014), 4.13 - Composition of the Oceanic Crust, in Treatise on Geochemistry (Second Edition), edited by H. D. Holland and K. K. Turekian, pp. 457-496, Elsevier, Oxford.

Yogodzinski, G.M., Rubenstone, J.L., Kay S.M., Kay, R.W., 1993, Magmatic and tectonic development of the western Aleutians: an oceanic arc in a strike-slip environment: Journal of Geophysical Research, v. 98. no. 7, p. 11807-11834.

Yogodzinski, G.M., 1993, Processes and components contributing to the formation of arc volcanic rocks: evidence from the western Aleutians [Ph.D. Thesis]: Ithaca, New York, Cornell University, 256 p.

Yogodozinski, G.M., Brown, S.T., Kelemen, P.B., Vervoort, J.D., Portnyagin, M., Sims, K.W.W., Hoernle, K., Jicha, B.R>, Werner, R., 2015, The Role of Subducted Basalt in the Source of Island Arc Magmas: Evidence from Seafloor Lavas of the Western Aleutians: Journal of Petrology, v. 56, p.441-492.

Yogodzinski, G.M., Kelemen, P.B., Hoernle, K., Brown, S.T., Bindeman, I., Vervoort, J.D., Sims, K.W.W., Portnyagin, M., Werner, R., 2017, Sr and O isotopes in western Aleutian seafloor lavas: Implications for the source of fluids and trace element character of arc volcanic rocks: Earth and Planetary Science Letters, v. 475, p. 169-180.

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Zimmer, M.M., Plank, T., Hauri, E.H., Yogodzinski, G.M., Stelling, P., Larsen, J., Singer, B., Jicha, B., Mandeville, C., Nye, C.J., 2010, The Role of Water in Generating the Calc-alkaline Trend: New Volatile Data for Aleutian Magmas and a New Tholeiitic Index: Journal of Petrology, v. 51, p. 2411-2444.

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Chapter 2: Rat (Hawadax) Island in the Context of the Aleutian Arc and Bowers Ridge

INTRODUCTION

Plate tectonics makes our planet fairly unique in the solar system and may be responsible for the conditions that allowed the formation of life and the continuing ability to maintain it. On the large scale, the surface of our planet is shaped by the motion of tectonic plates; however, to truly understand the details of the Earth’s structure, we must understand the causes of smaller scale features such as subduction zones and their associated island arcs, and on a smaller scale still, the placement, timing, and composition of volcanoes and individual islands, and more importantly, their evolution through time.

Rat (Hawadax) Island is a small, 13 x 3 km, island situated in the middle of the Rat Block in the central Aleutian Island Arc where Bowers Ridge (Ludwig et al., 1971) meets the Aleutian

Ridge at ~178° E (Figure 2.1). The focus of this study is to determine the geochemical and tectonic placement of Rat Island in the complex Bering Sea region. Because of its placement on the Aleutian Ridge at the terminus of Bowers Ridge, questions regarding the development and evolution of Rat Island arise. Is it a terminal part of Bowers Ridge or a volcanic island on the

Aleutian Ridge? Does its geochemistry represent an Aleutian Arc signature, a Bowers Ridge signature, or a combination of the two (perhaps dominantly one and contaminated by the other)?

Where does its age and geochemistry place it in the tectonic evolution of the Bering Sea region?

Can it be placed in the geochemical and tectonic history of the arc it belongs to (i.e., changes through time and along the length and width of the arc), or does its location give it unique characteristics? In addition, Rat Island and the Rat Block are located just west of the

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Figure 2.1: A) Map showing the location of the Bowers and Shirshov Ridges in relation to the Aleutian Arc and the locations of Attu, Adak, and the Rat Island Group along the Aleutian Arc. B) Area of box in Figure 1A. Close up of relevant islands and Bowers Ridge (west to east: Attu, Buldir, Kiska, Segula, Rat, Little Sitkin, Amchitka, Semisopochnoi, Adak, and Kagalaska). Maps produced using GeoMapApp (www.geomapapp.org).

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Aleutian subduction zone’s southernmost point in an area where the convergence rate between

the Pacific and North American Plates rapidly decreases (Syracuse and Abers, 2006). To the

degree that the character of subduction zone volcanism depends on convergence rate, we would

expect rapid changes in the former in this area. To this end, we look at the evolution of Rat

Island through time as well as its place in the evolution of the Rat Island Block.

Previously, little information on Rat Island was available; and we could only speculate as to the answers to these questions. This study, supported by new geochemical and

geochronological data, recent studies of Bowers Ridge, and several tectonic models, investigates

these issues and attempts to place Rat Island in the complex history of the Bering Sea region.

Clarification of Rat Island’s placement within the geochemical and tectonic context of the

Aleutian Arc and the Bering Sea region has the wider implication of enhancing our knowledge of

the effects of differences in the tectonic setting of subduction zones and their effects on the

geochemistry and evolutionary path of the resultant arc.

Bowers Ridge is a bathometric high, which is located north of the Aleutian Ridge in the

Bering Sea. A seamount chain that extends off the western tip of Bowers Ridge stops 10 km

short of connecting with Shirshov Ridge, indicating different structural origins (Rabinowitz,

1974). Keinle (1971) interpreted the end of the Rat-Near Island aftershock zone (1961-1967

earthquakes), where the Aleutian Ridge meets the seismically inactive Bowers Ridge, to indicate

an arc-arc transform between Bowers Ridge and the Aleutian Ridge. They concluded that the

Aleutian and Bowers Ridges are separate structural features, at least in the present and likely in

the past. Though the origin of Bowers Ridge was debated for many years, many believed it to be

a volcanic island arc that was produced by a now-inactive subduction zone, which either formed

in place or was transported to its current position by plate motions (Keinle, 1971; Cooper et al.,

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1981; Ben-Avraham and Cooper, 1981; Cooper et al., 1992; Scholl, 2007; Wanke et al., 2012;

Sato et al., 2016). Recent studies confirm the island arc nature of Bowers Ridge, though the

seamount chain on the western end displays a more MORB-like signature (Wanke et al., 2012;

Sato et al., 2016). Sato et al. (2016) also conclude that given the age and geochemistry of Bowers

Ridge, it formed in place as a SW-dipping subduction zone.

Based on the structure of the 8-10 km sediment wedge north of the ridge, Cooper et al.

(1981) originally placed the age of Bowers Ridge subduction as Mesozoic to Early Tertiary

(older than the Aleutian Ridge) possibly ceasing with Aleutian Ridge formation, but definitely inactive as the ridge subsided below sea level by the Middle Miocene. In the model of Ben-

Avraham and Cooper (1981), Bowers Ridge was formed to the west by subduction due to compression produced by a spreading ridge to the east of Bowers Ridge in the Bering Sea. Then,

Bowers Ridge migrated along an E-W trending transform, with subduction along the Beringian

Margin accommodating westward motion and subducting the spreading ridge. The collision of

the also westward-traveling Umnak Plateau with the Beringian Margin then caused subduction to

move south forming the Aleutian Ridge. Cooper et al. (1992) propose that Bowers Ridge was

once part of the Shirshov Ridge, but it moved SE along a backarc shear zone to its present

location. Extension due to the 42 Ma plate reorganization created Bowers Basin by back arc

spreading, with the migration of Bowers Ridge accommodating this spreading motion. In this

model, Bowers Basin formed after Bowers Ridge and continued spreading until ridge activity

ceased. New evidence suggests that Pacific Plate motion changed at > 50 Ma, possibly due to the

initiation of the Izu-Bonin-Mariana and Tonga Subduction zones (O’Connor et al., 2015).

The model of Yogodzinski et al. (1993) proposes that Bowers Ridge was inactive by ~43

Ma, but acknowledges that it may have been active until the Middle Tertiary as proposed by

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Cooper et al. (1992), but it is now known that Bowers Ridge was active as late as 26-22 Ma, as

discussed below. In Yogodzinski et al.’s model, at ~25 Ma, there was westward migration of arc- related blocks that eventually accreted onto the Kamchatka Margin as well as transportation of the Komandorsky Islands to their present location. Rostovtseva and Shapiro (1998) also support this transportation of the Komandorsky Islands based on the composition of their sedimentary rocks. At ~15 Ma, these collisions resulted in a change from transtensional to transpressional tectonics, resulting in a change from voluminous tholeiitic magmatism to less voluminous calc- alkaline magmatism in the western and central Aleutians.

The model of Scholl (2007) has Bowers Ridge forming in situ at approximately 45 Ma with subduction from the NE accommodating the movement of the orogenic stream (defined by

Scholl as a broad counterclockwise curving arc of laterally moving crust that extends from NW

British Columbia, through central Alaska, and eventually to the Kamchatka-Aleutian junction) across the Beringian Margin continuing through the Late Oligocene or Early Miocene. Scholl’s

(2007) model takes into account more tectonic parameters and places the timing of Bowers

Ridge subduction during or slightly after the formation of the Aleutian Ridge. This is supported in part by Wanke et al. (2012), who reported that dredged igneous rocks from the northern

Bowers Ridge were formed between 32 and 26 Ma and have trace element patterns and Sr-Nd-

Pb isotope compositions similar to the Western Aleutian Arc, implying formation above a subduction zone. Lavas from Bowers Seamount, located west of the Ridge proper, are younger

(24-22 Ma) alkali basalts with MORB-like isotopic compositions. Wänke et al. (2012) interpreted these as small degree decompression melts of a subduction-modified mantle. Given the relatively recent end of subduction, the timing of Scholl’s (2007) model seems quite reasonable.

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This study investigates this timing in the larger context of the Aleutian Arc and its surroundings using geochemical and geochronological data. It examines the geochemistry of Rat

Island as compared to the younger volcanic centers (15-0 Ma) of the Eastern and Western

Aleutian Arc as well as Bowers Ridge. By comparing the geochemistry and the environment that produced it, we gain insight into the processes forming the Aleutian and Bowers Ridges at the time Rat Island formed and whether there is a geochemical link between the two. In this study, we utilize major and trace elements (K2O, FeO/MgO, Ba/La, Sm/Yb, La/Sm, Ta/La, and Th/La)

87 86 and isotopic ratios (εNd and Sr/ Sr) to look at the source characteristics (mantle chemistry, depth and percent of melting), magma transport (speed, contamination, fractionation), contamination (by slab melts, slab fluids, and by the crust during transport), and eruption processes (location, petrology, contacts, and alteration) of Rat Island as compared to those of

Bowers Ridge and several young (<15 Ma) volcanoes and volcanic islands along the Aleutian

Arc, including Attu and Buldir to the west; the neighboring Rat Islands situated on the Rat

Block, Kiska, Segula, Little Sitkin, Amchitka, and Semisopochnoi; and Adagdak and Moffett to the east (Figure 2.1). Based on these geochemical comparisons, we can determine whether the chemistry of Rat Island is representative of the Aleutian Arc, Bowers Ridge, or a mixture of the two. This data can also be used to determine the nature, extent, and causes of any chemical changes (major element, trace element, and isotopic) that occur on Rat Island over time as well as the presence of any volcanic migration (i.e., away from the trench as also occurs elsewhere in the arc, such as on Adak, Kiska, and Attu). Through these investigations, we can gather valuable information that will allow us to investigate the tectonic and magmatic history of Rat Island in relation to the rest of the Aleutian Arc as well as the Bering Sea region.

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Figure 2.2: Map of Rat Island depicting sample locations and geologic units (after Lewis et al., 1960 – plate 70).

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Newly collected samples from the northern shore of Rat Island (Figure 2.2) were

analyzed for major and trace elements, Sr and Nd isotopes, and 40Ar/39Ar dating. This new data combined with the data of Wanke et al. (2012) from Bowers Ridge and Bowers Seamount allow for the geochemical comparison of Rat Island with Bowers Ridge. This comparison will provide the framework for interpretation of the geologic history of the area of the Aleutian Ridge directly south of Bowers Ridge.

GEOLOGIC BACKGROUND

Rat Island

The only previous study of Rat Island (Figure 2.2) is the USGS survey (Bulletin 1028) of

Lewis et al. (1960). They describe the island as 9.6 km long by 3.2 km wide with both subaerial

and submarine volcanism. High angle normal faults on the island have two trends: N60°W

(parallel to the Aleutian ridge) and approximately north. The presence of many linear elements in

the topography caused the authors to infer the presence of additional faults that make up a shear

zone. The higher areas show signs of glaciation, but glaciation is not noticeable in the lower

areas. Due to a lack of moraines, Lewis et al. (1960) infered that the entire island was glaciated.

The submarine bathymetry of Rat Island is fairly similar to that of Kiska and Amchitka,

indicating similar histories for these areas. To the southwest and possibly the northwest at a

depth of 90 m is a bench, to the northeast there is a steep slope, and to the south at a depth of 50

m is another bench that transitions to a steep slope. There is another shallower bench at a depth

of 15 m around almost the entire island. These benches likely represent changes in sea level

preserved by wave-cutting. Lewis et al. (1960) inferred a history beginning with the subaerial

eruption of porphyritic andesite and breccia flows interbedded with marine conglomerate (before

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the Rat Formation), faulting and chemical alteration, erosion and subsidence, mostly submarine basaltic volcanism in the Mid-Tertiary, subsidence (thought to also have occurred on Amchitka and southern Kiska), and finally uplifting, glaciation, and erosion.

Lewis et al. (1960) divided the exposed rocks of Rat Island into two formations (Figure

2.2): the Rat and Gunners Cove Formations. The contact between them was considered to be a

N-S trending high angle normal fault. The age of the Rat Formation, which Lewis et al. (1960) interpreted as being older than the Gunners Cove Formation based mainly on inclusions of what appears to be Rat Formation andesites in the Gunners Cove Formation, was only constrained as

Tertiary or older. The Rat Formation, which covers the southeastern third of the island, was described as being similar to the Amchitka Formation on Amchitka Island. As described by

Powers et al. (1960), the Amchitka Formation is composed of “water-laid beds of volcanic breccia and some thin-bedded fine grained tuff, all interbedded with pillow flows of lava.” The volcanic rocks are largely andesitic and usually porphyritic with phenocrysts of plagioclase, pyroxene, and amphibole. Their age was argued to be as old as Early Tertiary inasmuch as the overlying Banjo Point rocks on Amchitka are of Early Middle Tertiary age.” The volcanic units of the Rat Formation include thick columnar jointed flows and massive flow breccias of porphyritic hornblende andesite along with the conglomerates derived from them. The porphyritic andesite often contains large, zoned plagioclase (1 cm). Porphyries such as those seen in the Rat Formation are exceedingly rare in the Aleutian Arc. A special feature on Rat

Island is a zone of intense alteration rich in pyrite located approximately 4 km SE of Gunners

Cove.

Lewis et al. (1960) described the Gunners Cove Formation, which covers the northwestern two thirds of the island, as similar to the Amchitka Is. Banjo Point Formation.

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Powers et al. (1960), describe the Banjo Point Formation as “Bedded marine sandstone, conglomerate, tuffaceous shale, and some lapilli tuff of basaltic composition. The Banjo Point formation is placed, on the basis of a meager fauna, in the range Oligocene to Miocene.” Based on crinoid and pectin fossils Lewis et al. (1960) also estimated the age of the Gunners Cove

Formation to be Oligocene or Miocene. In order of abundance, this formation is composed of sandstone and conglomerate derived from basaltic volcanic rocks, some porphyritic hornblende andesite (likely weathered out of the Rat Formation), and shell fragments; basaltic tuffs; and thin basaltic flows, with pillow structures and poor columnar jointing, in thick scoria beds, intruded by dike swarms.

Table 2.1: Rat Island 40Ar/39Ar Ages Material Age (Ma) Formation RAT -13-1 groundmass 15.18 ± 0.86 Dike in Gunners Cove RAT-13-5 groundmass 13.35 ± 0.28 Rat Formation RAT-13-6 groundmass 14.33 ± 0.39 Rat Formation

RAT-13-8 plagioclase 12.91 ± 0.05 Rat Formation Kay et al., 2014

New 40Ar/39Ar ages determined by Brian Jicha of the University of Wisconsin are the first obtained for Rat Island. A Gunners Cove basaltic dike from the extreme north of the island

(Figure 2.2) was dated at 15.18 Ma ±0.86 (Table 2.1). As this dike cross-cuts older Gunners

Cove volcanics, it provides a minimum age for Gunners Cove magmatism. Attempts to date other Gunners Cove samples, which are typically highly altered, were unsuccessful. However, the age relationships are fairly ambiguous as the clasts in the Gunners Cove Formation (reported by Lewis et al., 1960) indicate that the Rat Formation should be older. As the dike’s age is within error of the oldest Rat Formation age, it is possible that the two formations are contemporaneous

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or very close in age. Three plagioclase-amphibole porphyritic andesites of the Rat Formation yielded ages ranging from 14.3 to 12.9 Ma. These dates correlate well with the pulse of magmatic activity that occurred elsewhere along the arc at 11-16 Ma (Scholl et al., 1976; Kay et al., 1982, 1994; Citron et al., 1980; Fournelle et al., 1994; Jicha et al., 2006).

Bowers Ridge

Until the geochemical studies of Wanke et al. (2012) and Sato et al. (2016), knowledge of

Bowers Ridge was restricted to geophysical data and interpretations. The ~700 km long Bowers

Ridge decreases in elevation as it curves counterclockwise away from the Aleutian Ridge and ends in an approximately E-W trending chain of seamounts (Ludwig et al., 1971; Scholl et al.,

1975). This seamount chain may be due to subduction from the NE resulting in more strike-slip subduction along the seamount chain (Keinle, 1971). The ridge itself is made up of three subaerially eroded plateaus separated by sediment filled depressions (Ludwig et al., 1971). Based on their geophysical investigations (two-ship seismic refraction measurements and closely spaced airgun-sonobuoy stations), Ludwig et al. (1971) describe the crustal lithologies of the ridge as volcanics and igneous intrusives similar to those of the Aleutian Ridge, with a crustal root extending to approximately 26 km bsl and a sediment filled trough parallel to the convex side of the ridge.

The geochemical study of Wanke et al. (2012) added a new dimension to knowledge of

Bowers Ridge. Dredge samples were taken from four sites on the northern slope of the ridge and one site on the seamount chain. The ridge samples have 40Ar/39Ar ages from 32.3±2 to 26 Ma,

and the seamount samples have ages from 24.4±0.8 to 22.2±2.7 Ma. Wanke et al. (2012)

describe Bowers Ridge rocks as calc-alkaline, medium-K rocks with major elements within the

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range of Aleutian Arc rocks. They note that the Bowers Ridge samples have more depleted Nb

and Ta and more enriched light rare earth elements (LREE), Ba, U, and Sr than the Aleutian Arc.

In addition, they point out that the Bowers Ridge samples have LREE and mobile elements (Ba,

U, K, Pb) enriched over NMORB, primitive mantle-normalized negative high field strength elements (HFSE: Nb, Ta, Ti), and arc-like anomalies. Sr (87Sr/86Sr: 0.70296-0.70311) and Pb

(206Pb/204Pb: 18.22-18.30; 207 Pb/204Pb: 15.45-15.46; 208 Pb/204Pb: 37. 65-37.73) are relatively

unradiogenic with radiogenic εNd (9.46-9.75) that falls between the central and western Aleutian

values. They describe the seamount samples as having moderately enriched LREE, moderately

fractionated HFSE; elevated large ion lithophile elements (LILE: Rb, K, Ba, Sr), Pb, Th; and

LREE concentrations that show a moderate subduction signature, and a Pacific MORB isotopic

signature. The seamount average LREE falls in the upper part of the Ridge samples enrichment

range. (Wanke et al., 2012)

These characteristics are interpreted by Wanke et al. (2012) to be the result of orthogonal

subduction at the central ridge and more oblique subduction along the seamount chain.

According to them, this oblique subduction and/or a slab tear account for the higher than normal

heating needed to melt the eclogitized seawater-altered slab and produce the observed trace

element and isotopic signatures. Small degree decompression melts of mantle upwelling from

Bowers Basin spreading or along the edge of the slab tear is thought to result in the observed

seamount chemistry. (Wanke et al., 2012)

Sato et al. (2016) studied the geochemistry and geochronology of dredge samples taken

from Bowers Ridge. They conclude that Bowers Ridge is an island arc formed by a SW-dipping

subduction zone that formed in place at approximately the same time as the Aleutian Ridge with

two identifiable episodes of volcanism, 34-32 Ma and 28-26 Ma. They agree with Wanke et al.

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(2012) that the 24-22 Ma seamount chain is not subduction related and has a more MORB-like geochemical signature that only seems to be present on the western-most tip of the ridge.

Bowers Basin

Bowers Basin has a transitional seismic layer between the upper mantle and the upper 6.1 km/s layer, with a p-wave velocity of 7.3 km/s, which is lacking in the Aleutian Basin, indicating structural differences (Ludwig et al., 1971). The velocity cross section of Ludwig et al. (1971) shows that this 7.3 km/s section is ~10 km thick near the Aleutian Ridge, decreases to ~2 km thick in the middle of the Basin, and increases to ~10 km thick near Bowers Ridge, which they caution may indicate that this layer may reflect the velocity structure of the Ridge systems. In addition to this, Bowers Basin displays negative to positive gravity anomalies east to west; whereas, the majority of the Aleutian Basin is in isostatic equilibrium (Keinle, 1971). The

Aleutian Basin has a 7-9 km thick igneous crust overlain by 4-6 km of sediment (Ben-Avraham and Cooper, 1981); Bowers Basin has an igneous crustal thickness similar to that of the Aleutian

Basin (7-9 km) and 2 km of sediment (Cooper et al., 1981); and the Komandorsky Basin has a 4-

5 km thick igneous crust and 1-2 km of sediment (Ben-Avraham and Cooper, 1981). The

Aleutian, Bowers, and Komandorsky Basins have heat flows of 1-1.8 HFU (heat flow units)

(Ben-Avraham and Cooper, 1981), ≥ 2 HFU (Cooper et al., 1992), and 2-4 HFU (Ben-Avraham and Cooper, 1981), respectively. These values place the age of Bowers Basin between that of the

Late Mesozoic Aleutian Basin (Cooper et al., 1976; Cooper et al., 1992; Scholl, 2007) and the

20-9 Ma Komandorsky Basin (Creager et al., 1973; Valyasho et al., 1993; Scholl, 2007).

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Table 2.2: Rat Island Sample Descriptions Sample Longitude Latitude Phenocrysts Description RAT-13-1 178.22075°E 51.83382°N Aphyric basalt, dike complex in Gunners Cove Formation RAT-13-2 178.24246°E 51.82714°N Minor plag., basaltic andesite clast in olivine hyaloclastite in Gunners Cove Formation RAT-13-2A 178.24246°E 51.82714°N basalt clast of hyaloclastite in Gunners Cove Formation RAT-13-2B 178.24246°E 51.82714°N Amphibole, Rounded dacite clast (Rat plagioclase Formation) in Gunners Cove Formation from upstream RAT-13-2C 178.24246°E 51.82714°N Plag., amph, Rounded boulder, ~1 ft across, minor cpx Gunners Cove Formation RAT-13-2D 178.24246°E 51.82714°N Boulder fragments, basalt hyaloclastite, Gunners Cove Formation RAT-13-3 178.27719°E 51.82307°N Plagioclase, basaltic, fresh, curved dike minor cpx and cutting hyaloclastite in olivine Gunners Point Formation RAT-13-4 178.30386°E 51.81909°N Plagioclase, Fresh andesite dike, North of minor cpx and Gunners Cove, slightly inland olivine mostly from RAT-13-5 altered to iddingsite RAT-13-5 178.30882°E 51.81927°N Large plag. and Rat Formation andesite, amphibole outcrop on west side of beach, porphyry RAT-13-6 178.32243°E 51.80770°N Plag., heavily Rat Formation andesite, South altered amph., of Gunners Cove, from beach oxides; rims on boulders, hyalaclastite plagioclase porphyry RAT-13-7 178.33707°E 51.80342°N Plag., heavily Rat Formation andesite, altered altered amph., oxides, rims on plag porphyry RAT-13-8 178.35017°E 51.78866°N Plag., amph., Rat Formation andesite minor oxides, porphyry, South of alteration, rims on plag extensive alteration zone, from porphyry right of massive outcrop Note: See Appendix C for detailed thin section descriptions.

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Figure 2.3: Pictures of Rat Island showing a) Gunners Cove; b) alteration zone, near RAT-13-8; c) dike, near RAT-13-2; d) breccia contact, near RAT-13-6; and e) hornblende and plagioclase porphyry, near RAT-13-6. (Sue Kay, personal communication)

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151

Figure 2.3 (Continued):

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Figure 2.3 (Continued):

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SAMPLING AND ANALYTICAL METHODS

Samples from Rat Island were collected in August 2013, during a scientific cruise to the

western Aleutians aboard the US Fish and Wildlife vessel Tiglax by Suzanne Kay, Brian Jicha, and Allen Schaen (Figure 2.3). Sample descriptions and locations are presented in Table 2.2, and all of the locations are plotted on the map in Figure 2.2. Samples were analyzed for major elements by ICP-OES at Cornell University, for trace elements by ICP-MS at Colgate

University, and for Sr and Nd isotopes by TIMS at Cornell University. 40Ar/39Ar ages were provided by Brian Jicha, University of Wisconsin (Tibbetts et al., 2014 – AGU abstract).

Detailed analytical methods are described in Appendix A. Thin section descriptions for selected samples are presented in Appendix C.

RESULTS

Major Elements

Major element data for Rat Island is presented in Table 2.3. Basalts predominate among the Gunners Cove samples, although andesites are also present. In contrast, andesites predominate among the Rat Formation samples, with a single dacite. Although samples of both units fall within the medium-K field defined by Gill (1981) (Figure 2.4), as do most Aleutian Arc volcanics, the Gunners Cove samples are both richer in K2O and Na2O than the Rat Formation samples at a given SiO2 concentration. Bowers Ridge samples plot as medium-K trending to high-K at higher silica contents; while the Bowers seamount samples fall on the boundary between medium-K and high-K.

The Gunners Cove Formation is predominantly tholeiitic with the exception of RAT-13-1

from the dike complex on the northwestern tip of the island, while the Rat Formation is

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exclusively calk-alkaline (Figure 2.5). Bowers Seamount volcanics are tholeiitic to transitional,

while Bowers Ridge samples range from calc-alkaline to tholeiitic. The differences in the calc- alkaline and tholeiitic trends and their causes are discussed in detail in Chapter 1.

Table 2.3: Rat Island Major Elements RAT-13-1 RAT-13-2 RAT-13-2A RAT-13-2B RAT-13-2C RAT-13-2D Gunners Gunners Gunners Rat Gunners Gunners Cove Cove Cove Formation Cove Cove

SiO2 50.39 56.51 49.23 65.20 49.46 50.34

TiO2 0.68 1.04 0.94 0.46 0.98 1.00

Al2O3 16.81 16.02 17.69 15.58 17.79 18.01 FeO 7.83 9.52 10.89 4.09 11.55 11.78 MnO 0.13 0.21 0.18 0.06 0.18 0.19 MgO 9.17 2.31 5.47 2.69 4.89 4.89 CaO 9.75 6.07 10.23 6.20 8.89 8.98

Na2O 2.84 4.19 2.98 3.17 3.20 3.30

K2O 0.43 1.91 0.63 1.27 0.50 0.52

P2O5 0.12 0.51 0.22 0.11 0.25 0.26 Total 98.15 98.28 98.48 98.82 97.67 99.25

RAT-13-3 RAT-13-4 RAT-13-5 RAT-13-6 RAT-13-7 RAT-13-8 Gunners Gunners Rat Rat Rat Rat Cove Cove Formation Formation Formation Formation

SiO2 51.42 57.81 58.82 60.91 59.12 57.01

TiO2 1.07 1.25 0.50 0.50 0.54 0.50

Al2O3 17.06 15.67 18.75 17.77 17.58 18.56 FeO 11.94 9.88 5.93 5.69 6.34 5.68 MnO 0.20 0.19 0.11 0.11 0.14 0.13 MgO 3.82 2.52 3.52 2.14 3.47 3.30 CaO 9.49 5.33 7.16 7.32 6.92 7.03

Na2O 3.41 4.78 3.95 3.66 2.34 3.81

K2O 1.13 1.95 0.94 1.14 1.22 0.93

P2O5 0.35 0.55 0.13 0.11 0.11 0.14 Total 99.90 99.92 99.80 99.35 97.77 97.09

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Figure 2.4: SiO2 vs. K2O discrimination diagrams for A) Rat Island: calcalkaline Rat Formation, tholeiitic Gunners Cove Formation; B) Bowers Ridge (red) (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics. Black lines delineate low, medium, and high- potassium groupings (after Gill, 1981). Data sources are as follows for these islands and volcanoes: Attu (This manuscript; Rubenstone, 1984; Yogodzinski et al., 1993); Buldir (Coats, 1953; Kay and Kay, 1985; Kay and Kay, 1994); Kiska (This manuscript; George et al., 2003); Little Sitkin (Snyder, 1959; Delong, 1974; White et al., 1984; George et al., 2003; Rhiannon et al., 2003; Yogodzinski et al., 2010); Segula (Delong, 1974; McCulloch and Perfitt, 1981; Nelson, 1959; George et al., 2003); Semisopochnoi (Delong, 1985);; Amchitka (Kay, 1980; USGS, 2008; Kay et al., 1986); Moffett (Coats, 1952; Conrad et al., 1983; Debari et al., 1987; Kay and Kay, 1985; Kay and Kay, 1994; Kay et al., 1982; Marsh, 1976; Myers et al., 1985; Walker, 1974; Kay et al., 1986; Yogodzinski et al., 2010); Adagdak ((Coats, 1952; Debari et al., 1987; Kay and Kay, 1988; Kay and Kay, 1985; Kay and Kay, 1994; Marsh, 1976; Myers et al., 1985; Von Drach et al., 1986); Seguam and Umnak (Singer et al., 2007).

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Figure 2.5: SiO2 vs. FeO/MgO Tholeiitic vs. Calc-Alkaline discrimination diagrams for A) Rat Island: Rat Formation, Gunners Cove Formation; B) Bowers Ridge (red) (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics Black line delineates Tholeiitic/Calc-alkaline boundary (after Miyashiro, 1974). Data sources as in Figure 2.4.

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Table 2.4: Rat Island Trace Element Concentrations (ppm) RAT-13-1 RAT-13-2 RAT-13-2A RAT-13-2B RAT-13-3 Sc 28.3 15.7 30.9 14.3 22.0 V 197 183 365 107 357 Cr 562 3.84 38.9 55.5 4.73 Co 36.6 18.9 33.9 13.9 31.5 Ni 172 4.73 28.2 33.0 8.81 Cu 85.5 150 167 31.0 232 Zn 62.3 114 81.8 28.2 106 Rb 7.57 56.7 9.18 16.3 17.7 Sr 379 488 545 530 571 Y 13.3 39.4 21.1 8.66 23.0 Zr 60.4 162 43.2 72.6 85.7 Nb 1.35 4.48 0.800 1.19 1.83 Ba 163 545 177 359 311 La 4.29 15.6 5.19 7.56 9.42 Ce 10.8 42.4 12.3 17.5 23.5 Pr 1.68 6.00 2.08 2.38 3.78 Nd 8.00 27.3 10.3 9.85 17.7 Sm 2.19 6.47 2.87 2.16 4.49 Eu 0.801 1.85 1.07 0.755 1.44 Gd 2.40 6.43 3.18 2.20 4.53 Tb 0.395 1.01 0.533 0.309 0.715 Dy 2.30 5.85 3.28 1.61 4.13 Ho 0.486 1.23 0.707 0.328 0.857 Er 1.34 3.43 1.94 0.909 2.35 Tm 0.202 0.529 0.289 0.133 0.353 Yb 1.25 3.28 1.72 0.831 2.16 Lu 0.200 0.508 0.273 0.131 0.331 Hf 1.69 4.55 1.30 2.06 2.4 Ta 0.089 0.291 0.063 0.106 0.124 Pb 1.39 7.39 2.23 3.540 4.01 Th 0.456 2.68 0.492 1.59 1.30 U 0.285 1.42 0.244 0.791 0.691

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Table 2.4 (Continued): RAT-13-4 RAT-13-5 RAT-13-6 RAT-13-8 Sc 20.0 14.1 13.6 11.7 V 189 115 143 114 Cr 4.69 32.4 49.6 30.8 Co 17.5 15.9 18.2 16.1 Ni 7.71 17.9 21.6 17.6 Cu 221 37.9 46.4 38.0 Zn 127 45.3 56.0 49.7 Rb 24.0 16.7 20.9 17.8 Sr 357 496 419. 513 Y 44.3 10.9 10.9 11.1 Zr 138 72.6 79.6 62.5 Nb 3.24 1.89 1.71 1.99 Ba 426 280 329 274 La 13.1 6.34 6.52 6.44 Ce 36.5 14.9 15.1 15.1 Pr 5.61 2.07 2.05 2.08 Nd 26.6 8.80 8.45 8.78 Sm 6.88 2.04 1.91 2.01 Eu 2.03 0.777 0.728 0.775 Gd 7.02 2.19 2.07 2.16 Tb 1.15 0.333 0.320 0.325 Dy 6.93 1.84 1.80 1.80 Ho 1.46 0.382 0.382 0.374 Er 4.03 1.07 1.10 1.05 Tm 0.624 0.163 0.171 0.159 Yb 3.82 1.03 1.10 1.00 Lu 0.591 0.165 0.178 0.161 Hf 4.10 1.95 2.13 1.74 Ta 0.228 0.150 0.146 0.146 Pb 5.68 3.36 3.74 3.53 Th 1.76 0.956 1.16 0.930 U 0.912 0.612 0.713 0.578

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Trace Elements

Trace element data for Rat Island is presented in Table 2.4. Both the tholeiitic Gunners

Cove Formation and the calc-alkaline Rat Formation are moderately light rare earth element

(LREE)-enriched (Figure 2.6). The Gunners Cove samples have higher and more variable overall rare earth element (REE) concentrations, but their REE patterns are less steep than those of the

Rat Formation, despite the fact that the Gunners Cove samples have higher overall silica concentrations than the Rat Formation samples. The latter have notably consistent REE patterns, which is consistent with their more uniform major element compositions. The one Rat Formation dacite, RAT-13-2B, is slightly LREE enriched and slightly HREE depleted relative to the Rat

Formation andesites.

The Rat and Gunners Cove Formations show typical subduction-related incompatible

element patterns with enrichments in Ba, U, K, Pb, and Sr and depletions in Th, Nb, and Ta

(Figure 2.7). Some Gunners Cove samples also have a slight depletion in Hf and Zr. Gunners

Cove concentrations increase with silica content, while the Sr enrichment and Zr and Hf

depletions decrease. Though we used a bomb dissolution method for the more silicic samples,

there is still a possibility that the Hf and Zr depletions are the result of the samples not being

completely dissolved. Again, there is little variation in the Rat Formation concentrations, and

they are lower than those of Gunners Cove despite higher silica contents.

Although the absolute abundances of the incompatible elements of the Rat and Gunners

Cove Formations overlap, on average, the Rat Formation is more enriched in incompatible

elements than the Gunners Cove Formation, and the Rat Fm. also has slightly higher La/Sm for

samples with similar Sm/Yb (Figure 2.8). All of the Rat Island samples plot within the young

volcanic Aleutian range for both of these ratios, but the Gunners Cove volcanics tend to plot at

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the low end of the range for both ratios. Low Sm/Yb appears to be a characteristic of Rat Is. as

well as the rest of the Rat Island Group, despite variable La/Sm. The Rat Formation samples,

however, have nearly constant La/Sm with variable Sm/Yb. In the Gunners Cove samples,

La/Sm appears to correlate with Sm/Yb, although the range in both ratios is quite limited. The

Bowers Ridge samples have distinctly higher Sm/Yb than the Rat Islands, while the Bowers

Seamount samples have both La/Sm and Sm/Yb within the range of the Rat Islands.

Depletion in Ta and Nb and enrichment in Ba relative to other incompatible elements has

long been recognized as a distinguishing feature of subduction-related basalts (e.g., Gill, 1981;

Kay, 1984). Here we use a Ta/La ratio of ~0.4 to distinguish arc-like (Ta/La <0.4) from MORB- like (Ta/La<0.4) magmas (average Pacific MORB Ta/La is 0.058±0.018; White and Klein,

2014). All of the Rat Island and Bowers Ridge samples have arc-like Ta/La, as do all of the

Aleutian samples, except one Kiska sample with Ta/La just below 0.4 (Figure 2.9). The Bowers

Ridge samples have the lowest Ta/La. The Gunners Cove data extends to lower Ta/La than the

Rat Formation. In contrast, the Bowers Seamount samples all have MORB-like Ta/La. The

Gunners Cove samples have slightly higher Nb/Ta than the Rat Formation and Bowers Ridge

samples (not shown).

The Ba/La of the Rat Formation samples is systematically higher than that of the Gunners

Cove Formation, with the latter falling in the lower part of the central and western Aleutian Arc

range, and with Gunners Cove being similar to Little Sitkin (Figure 2.9). The Rat Islands span

the entire range of Ba/La, with Semisopochnoi being one of the most enriched and Segula one of

the least enriched. The Bowers Ridge samples span the Ba/La range of Rat Island and the

Aleutian samples. The Bowers Seamount samples have slightly higher Ba/La than the Pacific

MORB field and are much lower than the Aleutian and Bowers Ridge samples. There is little

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correlation between Ba/La and Ta/La. The Ba/La ratios of the Bowers Ridge and Seamount samples are generally lower than those of the Rat Group and Rat Island in particular. The Rat

Formation samples have systematically higher Th/La than those of the Gunners Cove Formation

(Figure 2.10). All of the Rat Island samples and the Rat Island Group as a whole overlap the range defined by the Aleutian samples, with Attu overlapping with Rat Island and the Adak

(Adagdak and Moffett volcanos) samples having the highest values. The lowest Gunners Cove samples just reach the edge of the Pacific MORB field. In contrast to Ta/La, Th/La shows an overall correlation with Ba/La. However, this correlation is almost entirely due to the Adak Mt.

Moffett samples. In contrast, the ratios do not correlate within the Rat Group or in the Western

Aleutians samples.

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Figure 2.6: CI-Chondrite normalized rare earth element plots for A) Rat Island (tholeiitic Gunners Cove – 1, 2, 2A, 3, 4; light blue lines) (calc-alkaline Rat Formation – 2B, 5, 6, 8; dark blue lines); and B) Bowers Ridge (red field), Bowers Seamount average (pink) (Wanke et al., 2012), young Aleutian volcanics (Adak and Kagalaska: Sue Kay personal communication; Attu: this manuscript; Buldir: Earthchem database), and young Rat Island Group volcanics (Kiska: this manuscript; Amchitka: USGS database; other: Earthchem database). In order on the legend, the locations are: Attu, Kiska, Adagdak, Kagalaska, Semisopochnoi, Bowers seamount average, Buldir, Little Sitkin, Moffett, Segula, and Amchitka. Normalization values from Sun and McDonough (1989).

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Figure 2.7: Primitive Mantle normalized spider diagrams for A) Rat Island (tholeiitic Gunners Cove – 1, 2, 2A, 3, 4; light blue lines) (calc-alkaline Rat Formation – 2B, 5, 6, 8; dark blue lines); and B) Bowers Ridge (red field), bowers seamount average (pink) (Wanke et al., 2012), young Aleutian volcanics (Adak and Kagalaska: Sue Kay personal communication; Attu: this manuscript; Buldir: Earthchem database), and young Rat Island Group volcanics (Kiska: this manuscript; Amchitka: USGS database; other: Earthchem database). In order on the legend, the locations are: Attu, Kiska, Adagdak, , Kagalaska, Semisopochnoi, Bowers seamount average, Buldir, Little Sitkin, Moffett, Segula, and Amchitka. Normalization values from Sun and McDonough (1995).

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Figure 2.8: Sm/Yb vs. La/Sm normalized to C1 chondrite (McDonough and Sun, 1995) for A) Rat Island: calcalkaline Rat Formation, tholeiitic Gunners Cove Formation; B) Bowers Ridge (red) (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Data sources as in Figure 2.4.

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Figure 2.9: La/Ta vs. Ba/La for A) Rat Island: calcalkaline Rat Formation, tholeiitic Gunners Cove Formation; B) Bowers Ridge (red) (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Black line divides low La/Ta ocean ridge-like samples from high La/Ta arc-like samples. Data sources as in Figure 2.4.

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Figure 2.10: Th/La vs. Ba/La for A) Rat Island: calcalkaline Rat Formation, tholeiitic Gunners Cove Formation; B) Bowers Ridge (red) (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Data sources as in Figure 2.4.

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Isotope Ratios

Sr and Nd isotopic ratios are listed in Table 2.5. As Figure 2.11 shows, all of the Rat

Island samples, as well as the other Rat Group Islands, fall within the range of the other young

Aleutian volcanics plotted, the high εNd end of which overlaps the lower end of the MORB field.

87 86 The Aleutian and Rat Island samples have lower εNd and slightly higher Sr/ Sr than those of

Bowers Ridge. By comparison, the Bowers Seamount samples plot within the range of Pacific

MORB, albeit at the extreme end of that range (Figure 2.11).

εNd and Th/La form a trend from the low Th/La and high εNd Bowers Ridge and

Seamount, which overlap the MORB range, to the higher Th/La and lower εNd Rat Islands

(Figure 2.12). Th/La and εNd scatter about this same trend. Thus, Rat Island and the Rat Island

Group volcanics were derived from a less incompatible-depleted source than Bowers Ridge and

Seamount.

Table 2.5: Rat Island Isotopes 87 86 143 144 Formation Sr/ Sr 2σ Nd/ Nd 2σ εNd 2σ RAT -13-2 Gunners Cove 0.703345 ±8 0.512995 ±14 6.96 ±0.28 RAT-13-2A Gunners Cove 0.703152 ±14 0.513094 ±14 8.89 ±0.28 RAT-13-4 Gunners Cove 0.703158 ±8 0.513046 ±20 7.95 ±0.38 RAT-13-6 Rat Formation 0.703171 ±7 0.513018 ±22 7.40 ±0.42 RAT-13-8 Rat Formation 0.703176 ±11 0.513087 ±9 8.75 ±0.18

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87 86 Figure 2.11: Sr/ Sr vs. εNd for Rat Island, Bowers Ridge (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Data sources as in Figure 2.4.

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Figure 2.12: εNd vs. Th/La for Rat Island, Bowers Ridge (Wanke et al., 2012), and young Aleutian volcanics. Black symbols: western Aleutians; purple symbols: central Aleutians; and green symbols: young Rat Island Group volcanics. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Data sources as in Figure 2.4.

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DISCUSSION

In the following discussion, we consider the evidence supporting the assignment of Rat

Island to either the Aleutian Ridge or Bowers Ridge based on compositional contrasts between them; the relationship between Rat Island’s position on the Aleutian Ridge, its chemistry, and its age and that of neighboring volcanism on the Ridge; and the agreement between the timing of

the evolutionary changes on Rat Island and Scholl’s (2007) model for the evolution of the

Aleutian Arc and the Bering Sea region. Using comparisons between Rat Island, the central

Aleutian volcanics, and the Rat Island Group volcanics, we look at the evolution through time of

both the Rat Island block and the Central Aleutians (Delarof and Andreanof Blocks; from

Amatignak to Seguam).

Rat Island: Part of Bowers Ridge or the Aleutian Arc?

A significant amount of time elapsed between the youngest Bowers Ridge volcanism and

the oldest Rat Island volcanism. The youngest age for Bowers Ridge is 26 Ma and for Bowers

Seamount 22 Ma (Wanke et al., 2012). Rat Island’s oldest known age is 15.18 Ma (Jicha,

personal communication), which is ~ 7 Ma younger than the youngest known age of Bowers

volcanism. In addition to this, the youngest volcanism for Bowers Ridge occurred at Bowers

Seamount off the northwestern tip of Bowers Ridge; whereas Rat Island is located at the opposite

end, south of the base of Bowers Ridge. Neither the timing nor locations of the two volcanic

systems coincide.

Ta/La ratios decrease from the Rat Island samples to the Bowers Ridge samples (Figure

2.9). Bowers Ridge on average has a more pronounced Ta depletion (Figure 2.7), indicating a

difference in the subduction/melting environment. Bowers Ridge’s Ta-depletion forms an

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endmember to the Aleutian Arc volcanics, while the Rat Island samples fall within the more

depleted half of the Aleutian field. Thus, with respect to Ta/La, Rat Island more closely

resembles the Aleutian Arc than Bowers Ridge.

87 86 The Sr/ Sr and εNd values for Rat Island fall completely within the Aleutian field. In contrast, Bowers Ridge has εNd on the low end of the MORB field, similar to those of the Bowers

Seamounts, but with higher than MORB Sr values that fall in the low end of the Aleutian field

(Figure 2.11). This difference in εNd indicates different source characteristics for the Bowers and

Aleutian Ridges and suggests Rat Island magmatism has the same source as the Aleutian Arc.

The trace element enrichments (Ba, U, K, Pb, and Sr) and depletions (Th, Nb, and Ta) of

the Rat Island and Bowers Ridge volcanics both indicate that they formed in a subduction-related

environment (Figure 2.7). The Ba/La and Th/La ratios of the Rat Island volcanics fall within the

Aleutian field and outside the MORB field, while Bowers Ridge and Seamounts have MORB-

like Th/La, which further highlights Rat Island’s subduction characteristics. Other compositional

features also emphasize the similarity of Rat Island to the Aleutian Arc and their difference from

Bowers Ridge and Seamounts. The flatter REE patterns of Bowers Seamount and its depletion in

Pb, lack of Nb and Ta deletions, smaller K enrichment, and lower Ba all indicate a more MORB-

like environment, which is confirmed by the more MORB-like isotopic values (Figures 2.6, 2.7).

This is reasonable considering the seamount chain’s placement on the edge of the subduction

zone where there was likely a large strike-slip component.

Relative enrichment in hydrous fluid-mobile elements, such as Ba, is generally similar in

the Bowers Ridge and Rat Island volcanics (Figure 2.9), while fluid-immobile elements, such as

Th and Ta, are relatively less enriched in Bowers Ridge. In addition, Bowers Ridge has slightly

87 86 lower Sr/ Sr values than the Aleutian Arc and Rat Island and distinctly higher εNd (Figure

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2.11). These attributes suggest either an inherently more depleted source for the Bower Ridge, or more likely, that the Aleutian mantle source had a greater slab-derived melt component than that of the Bowers mantle source.

The Bowers Seamount samples have Ba/La ratios that fall within or near the North

Pacific MORB field (Figure 2.10), indicating minimal contribution from subduction-derived fluids. This is consistent with the seamount chain being produced in a transtensional strike-slip environment with ocean ridge-like upwelling of a MORB source contaminated by nearby subduction, as was suggested by Wanke et al. (2012).

Some of the Bowers Ridge volcanics show a greater enrichment in K2O at a given SiO2

(Figure 2.4) than the Rat Island volcanics, in contrast to most other incompatible elements. This contrast with other trace element and isotopic indicators of source enrichment combined with their reasonably high Ba/La and >20 Ma age suggests that the higher K of the Bowers Ridge samples is most likely due to seawater alteration, or less likely, a difference in the style of fractional crystallization.

The high Sm/Yb of the Bowers Ridge samples (Figure 2.8) suggests significant residual garnet in the source. Several factors could produce this. First, a lower overall extent of mantle melting leaving more garnet in the residue, but this should produce generally higher concentrations of all incompatible elements; whereas, Bowers Ridge is not incompatible element-enriched compared to the Aleutian Arc. Given how little is known about the Bower subduction zone, this possibility can only be speculated upon. A second possibility is a more basaltic, garnet-rich source. This could imply either partial melting of subducting oceanic crust or melting of lower arc crust carried into the subduction zone by subduction erosion. In contrast,

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Rat Island has some of the lowest Sm/Yb of the Aleutian samples, and thus, Rat Island clearly does not share this source characteristic with Bowers Ridge.

Magmatic Evolution on Rat Island

The Gunners Cove Formation is tholeiitic, while the Rat Formation is calc-alkaline. This represents a distinct change from larger volume (covers 2/3 of the island), basaltic, tholeiitic magmatism on Rat Island to smaller volume (covers 1/3 of the island), andesitic, calc-alkaline magmatism. Elsewhere in the Rat Island Block, both tholeiitic (Semisopochnoi, Segula,) and calc-alkaline volcanic centers (Amchitka, Little Sitkin) occur. The older southern part of Kiska is tholeiitic, while the modern volcanism, confined to the northern end of the island, is calc- alkaline. As discussed in Chapter 1, the difference between calc-alkaline and tholeiitic trends can be related to early crystallization of oxide phases, and hence oxygen fugacity and water content, and delayed crystallization of plagioclase, and hence water content or the pressure of crystallization. The question presented here is which of these was responsible for the shift from tholeiitic to calc-alkaline volcanism on Rat Island?

On Rat Island and in other parts of the Aleutian Arc (e.g., Kay and Kay, 1994), including neighboring Kiska, the transition is from early more voluminous tholeiitic magmatism to later less voluminous calc-alkaline magmatism, so the mechanisms discussed in Chapter 1 are certainly possibilities in the case of Rat Island.

The Gunners Cove samples have somewhat lower Ba/La (Figure 2.10) than the Rat

Formation. As Ba is very fluid-mobile and La less so, this suggests a lower slab-derived hydrous fluid contribution, and therefore, less slab-derived water in the Gunners Cove volcanics, which is consistent with the study of Brounce et al. (2015). While the lower Ba/La could also be due to

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87 86 decreased sediment contribution, this would also influence the Sr/ Sr and εNd values, but these are similar for the two formations. Thus, the increase in Ba/La from Gunners Cove to the Rat

Formation suggests an increase in water content is responsible for the shift from tholeiitic to calc-alkaline volcanism. It should be noted that the Rat Formation samples are more evolved and have large amphibole and feldspar phenocrysts, which could increase the Ba content of these rocks as Ba is compatible in these minerals. This transition from tholeiitic to calc-alkaline at ~15

Ma may correspond to a change to a more compressive environment due to the clockwise block rotation that occurred at that time (55°±8° on Amchitka; Minyuk and Stone, 2009). Although thickening crust and increasing depth of crystallization might also play a role, these are less likely given the lack of time between the emplacement of the two formations.

Kay et al. (1982) observed that the least differentiated members of the Aleutian calc- alkaline and tholeiitic series are similar, implying a common mantle-derived parent, but the petrology of the lavas of the different series suggests that the tholeiitic series crystallized under lower pressure. They also observed that tholeiitic volcanoes occur primarily at the end of or between arc segments, while calc-alkaline volcanoes occur in the center of arc segments. They argued that compressive structural forces in the center of segments resulted in greater depths of crystallization, and thus, were responsible for the calc-alkaline differentiation trend. Rat Island is located at the center of the Rat Block, and hence, this hypothesis predicts that it should be calc- alkaline, as it was in the younger Rat Formation phase, but not in the early Gunners Cove phase.

Since the tectonic configuration at 14 Ma is not known, whether a change in tectonic stresses, as proposed by Kay et al. (1982), also played a role in the transition from tholeiitic to calc-alkaline on Rat Island can only be speculated upon. However, it is noteworthy that the change from tholeiitic to calc-alkaline magmatism on Rat Island does correspond with the ~15 Ma change

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Figure 2.13: Comparison of the Gunners Cove and Rat Formations: a) Sm vs. SiO2, b) La vs. SiO2, c) Yb vs. SiO2, and d) (La/Sm)N vs. SiO2. La/Sm in d is normalized to C1 chondrite (McDonough and Sun, 1995).

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Figure 2.14: La vs. La/Sm for the Gunners Cove and Rat Formations.

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from transtensional to transpressional tectonics in the model of Yogodzinski et al. (1993) and the

55° clockwise rotation of the Rat Block as seen on Amchitka (Minyuk and Stone, 2009).

Excluding the dacite (RAT-13-2B), the Rat Island Group and Bowers Seamount

volcanics exhibit variable (La/Sm)N (~1 to 2.8) with roughly constant (Sm/Yb)N of ~2±0.2

(Figure 2.8). However, (La/Sm)N is significantly higher in the calc-alkaline Rat Formation

volcanics (~2) than in the Gunners Cove volcanics (~1 to 1.5). Furthermore, rare earth patterns

and rare earth abundances in the Rat Formation volcanics are remarkably constant, possibly due

to the fact that they represent a fairly small range of silica contents and are from only a few

different locations. Whereas, both the steepness of the patterns and the overall concentration

levels are quite variable in the Gunners Cove volcanics, possibly due to a larger range of silica

contents. Rare earth concentrations correlate with SiO2 in the Gunners Cove volcanics, and

(La/Sm)N also appears to correlate, at least weakly, with SiO2 (Figures 2.13, 2.14). This suggests

that fractional crystallization (particularly of clinopyroxene, which strongly fractionates the light

rare earths) is controlling the rare earth abundances and patterns in the Gunners Cove Formation.

Rare earth abundances are effectively independent of SiO2 in the Rat Formation, although

(La/Sm)N does appear to increase slightly with SiO2 (Figures 2.13, 2.14), as is expected with

amphibole fractionation. However, there are too few data points to determine if this correlation is

significant. The consistency of rare earth abundances and patterns in the Rat Formation is likely

due to (1) fractionation of amphibole rather than clinopyroxene, as partition coefficients are

generally somewhat higher in the former; (2) higher mineral melt partition coefficients

associated with the higher SiO2 concentrations (e.g., Nash and Crecraft, 1985); and (3) buffering

by accessory phases, apatite in particular, which strongly concentrate rare earths, particularly the

lighter rare earths (Prowatke and Klemme, 2006), resulting in a plateau in REE enrichment and a

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lack of further enrichment with increasing silica content and fractionation. The higher Sm/Yb in

the dacite, RAT-13-2B, may be due to zircon fractionation. Zircon strongly concentrates the

heavy rare earths (e.g., Watson, 1980), and the Yb concentration in RAT-13-2B is lower than that of all of the other Rat Island samples.

The difference in the La/Sm ratios of the Gunners Cove and Rat Formations is most

likely due to a combination of a slightly smaller extent of partial melting and fractionation of

amphibole in the Rat Formation, which may be consistent with their calc-alkaline character. The positive Eu anomalies in the Rat Formation samples are due to the accumulation of plagioclase, and the overall pattern exhibits a classic amphibole loss shape, which is supported by the fact that these samples are amphibole/feldspar porphyries. These porphyries also most likely represent slow cooling and a relatively long residence time in the magma chamber. Fractional crystallization by itself can be ruled out, as this would also result in overall higher concentrations, which indicates that the extent of melting must have also played a role.

Differences in source enrichment can be ruled out as there is no systematic difference in εNd

between the two formations.

The Rat Formation has slightly lower Nb/Ta than the Gunners Cove Formation. As

Yogodzinski et al. (2015) noted, Nb/Ta is typically lower in the Aleutian calc-alkaline lavas than

in the tholeiitic ones and is lowest in the most silica-rich lavas. Amphibole has higher partition

coefficients for Nb than Ta. Thus, this may be due to amphibole fractionation as the Rat

volcanics are amphibole-rich, while amphibole is largely absent from the Gunners Cove

Formation. This also correlates with the less sodic, lower water content of tholeiitic rocks

(Gunners Cove) leading to little or no amphibole formation, in contrast to the larger amount of

amphibole fractionation in the more sodic, higher water content calc-alkaline rocks (Rat

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Formation) (e.g., Sisson and Grove, 1993). The Gunners Cove samples have lower Sr

concentrations than those of the Rat Formation at comparable SiO2. This is likely due to the early

fractionation of plagioclase in the tholeiitic trend compared to the calc-alkaline one. The

depletions in Hf and Zr at higher silica contents, i.e., andesite, in the Gunners Cove Formation is possibly due to the fractionation of one or more accessory phases such as rutile, ilmenite, or zircon.

Evolution of the Rat Island Block Through Time

Volcanism all along the Aleutian Arc has migrated northward, i.e., the youngest

volcanism is on the northern part of the islands (e.g., Kay and Kay, 1994). This migration can

also be seen as a trend from island to island within the Rat Island Group. Little Sitkin,

Semisopochnoi, Segula, and northern Kiska are the northern most parts of this group and all have

Holocene volcanoes. In addition, Little Sitkin, Semisopochnoi, and Northern Kiska have erupted

historically. There is no active volcanism on Rat and Amchitka, which are the islands closest to

the trench. This northward migration could be due to subduction erosion or shallowing of the

slab, both of which would require melting to take place farther away from the trench (north) to

maintain the same depth of melting.

The Rat Islands are interesting in that they do not conform to the normal single line of

islands, roughly parallel to the ridge axis, but are instead arranged in a group at various distances

from the trench. The question then becomes whether there is a correlation between their

chemistry and their distance from the trench. In order of increasing distance from the trench, the

islands are Amchitka, Kiska and Rat, Little Sitkin and Segula, and Semisopochnoi. Other

possibilities for trends are from east to west across the block or the edges of the block versus the

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middle of the block. It is unfortunate that there is little age data for most of the volcanic centers in the Rat Island block, so a correlation with time is not possible. However, based on relative stratigraphic and weathering age constraints, it is clear that the younger volcanism on all of these islands is to the north. The exception to this is that on Rat Island, the older Gunners Cove

Formation covers the northwestern 2/3 of the island and the younger Rat Formation covers the southeastern 1/3 of the island. However, it should be noted that the dates for Rat Island indicate both formations were erupted within a fairly short timeframe, and the currently active Little

Sitkin is located directly to the north of Rat Island. So although the migration is not seen within the relatively small area of Rat Island, it can be seen among the islands of that part of the Rat

Block.

No trends are seen from east to west or for the middle versus the edge of the block

(Figure 15). There are also no apparent trends in chemistry across the width of the arc, i.e., with distance from the trench within the Rat Island Block (Figure 2.12), although there is little or no data for some of these ratios on some islands. This suggests that distance from the trench is not the controlling factor in the variation in these ratios.

The one tholeiitic Amchitka sample is quite old (33 Ma at least); the other samples, which are all calc-alkaline, are 33 Ma or younger, with two samples that are 5 Ma at most (Kay,

1980; USGS, 2008; Kay et al., 1986). The calc-alkaline Little Sitkin samples are believed to be

2.6 Ma at most (Snyder, 1959; Delong, 1974; White et al., 1984; George et al., 2003; Rhiannon et al., 2003; Yogodzinski et al., 2010). Unfortunately, the best age constraints for the

Semisopochnoi tholeiitic samples are “Late Pleistocene” based on evidence of glaciation (Coats,

1959), with the entire sequence believed to have been “erupted over a short enough period of time to have shared the same system of magma conduits and chamber(s)” by Delong et al.

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(1985). There is also a historic, 1987, eruption on Semisopochnoi (Smithsonian Global

Volcanism Program), indicating that volcanism is ongoing. The best age constraint for the

Segula tholeiitic samples is that they are believed to be Late Quaternary in age based on the lack

of glaciation and the small amount of marine erosion (Nelson, 1959). Kiska transitions from

tholeiitic to calc-alkaline at ~13 Ma. While Rat Island transitions form tholeiitic to calc-alkaline at ~15 Ma. The lack of age constraints on the islands that have samples in both fields makes it hard to determine if the other islands in the Rat Island block follow the transition from tholeiitic to calc-alkaline that the rest of the arc (including Rat, Kiska, and Amchitka) follows. However, from the little information we have, it seems likely that the rest of the Rat Island Block follows the pattern of the rest of the Aleutian Arc regardless of their distance from the trench.

A recent study by Schaen et al. (2016) on the Delarof Block, the next block to the east of the Rat Island Block, investigates this trend in northward migration with time from Amatignak and Ulak (37-27 Ma) in the southern Delarof Islands to Kavalga, Ogliuga, and Skagul (4.8-6.3

Ma) in the central Delarof Islands to Gareloi and Tanaga (Pleistocene-present) to the north. This study included 37 Ma to historic volcanism. Over this longer time frame, it was possible to see, not only the northward migration of volcanism with time, but also a trend toward more sediment and fluid input over time. They concluded that this change in chemistry (also seen on Adak; Kay and Kay, 1994) was due to the change in plate motion at 6 Ma, which caused block rotation and increased fragmentation and orthogonality of subduction, and was also possibly due to the increase in sediment input to the trench from glaciation beginning at ~2.7 Ma. It is possible that in a study of the Rat Island Block over its full geologic history, we would see a similar trend; however, within this study’s shorter time frame (due to the focus on Rat Island, which only has younger ages) we did not see this chemical shift. It would also not be possible to explore this

183

87 86 Figure 2.15 a) (Sm/Yb)N, b) (La/Sm)N, c) Ba/La, d) Ta/La, e) Th/La, f) Sr/ Sr, and g) εNd. Legend is in order from closest to the trench (left) to farthest from the trench (right). Amchitka, Rat and Kiska, Little Sitkin and Segula, Semisopochnoi. Sm/Yb and La/Sm are normalized to C1 chondrite (McDonough and Sun, 1995). Data sources as in Figure 2.4.

184

185

Figure 2.15 (Continued):

186

trend back in time for Rat Island by looking at islands further south (as was done in the Delarof

Block) due to the fact that there are no islands directly south of Rat Island. The transition in

chemistry at ~6 Ma (no 6 Ma samples on Rat Island) could potentially be investigated by looking

to the north to Little Sitkin, which is very similar in chemistry to Rat Island. This would indicate

that the ~6 Ma chemical transition seen on the two blocks to the east (Delarof and Andreanof

Blocks) and in the Near Island block (Attu) is not seen in the Rat Island Group. It should be

noted that the 33 Ma Amchitka sample had the highest εNd in the Rat Island Block and is the

oldest sample plotted. This follows the trend seen in the Delarof Block, which included older

samples that were closer to the trench.

Overall, it seems that the Rat Block follows the pattern of northward migration of

volcanism that is observed elsewhere in the Aleutians. The evolution with time affects all of the

islands in approximately the same way rather than varying with location. The cause or causes of

the northward shift in volcanism and the changes in chemistry with time appears to be

widespread, and not only affects the entire length of the arc, but the entire width of the arc as

well. Changes in the geometry of subduction could change the compressive forces, depth of

melting, input of sediments and fluids, and the amount of subduction erosion. These changes

would cause shifts in the composition and location of magma production along the width and

breadth of the arc. For example, a change in fluid input at one point in the arc would change the

fluid enrichment of the mantle wedge at all distances from the trench though not necessarily to the same degree or simultaneously, i.e., areas farther from the slab would take longer to be effected. However, these changes would not necessarily be identical along the arc, as there are

significant differences in several controlling factors along arc, e.g., sediment supply, subduction

187

angle and rate, slab dip, depth of the seismic zone, etc. However, the distance of the islands from the trench does not display any linear correlation with chemistry.

This lack of an apparent chemical change with distance from the trench could indicate that the distance from the trench was approximately the same for these eruptions. That is, as volcanism was moving northward, so was the trench. If subduction was eroding the overlying plate and shifting subduction northward with time, and with it the locus of volcanism, the volcanic centers could have been approximately the same distance from the trench at the time they erupted, which would explain why we do not see a trend in chemistry with distance from the trench. This has been seen in other subduction zones, such as the Andes (Goss et al., 2013;

Scholl and von Huene, 2009; Stern, 2011; Stern and Scholl, 2010). If subduction erosion was mostly transporting volcanic material with an isotopic signature similar to that of the erupting volcanics into the trench, we wouldn’t necessarily see a distinctive geochemical change to mark the presence of the subduction erosion.

CONCLUSIONS

1) Rat Island is part of the Aleutian Arc and has no relationship to Bowers Ridge. This is supported by the following observations:

• The oldest volcanism on Rat Island is 10 Ma and 6 Ma younger than that on Bowers

Ridge and Bowers Seamounts, respectively, and occurs on the end of the ridge opposite

Rat Island, indicating that Rat Island magmatism occurred after magmatism on Bowers

Ridge ceased.

• Bowers Ridge is more depleted in fluid-immobile elements such as Ta and Th and has

more depleted Sr and Nd isotopic values than the Aleutian Arc and Rat Island, which

188

falls within the Aleutian field, indicating that Bowers Ridge has a more depleted mantle

source and less fluid input from subduction.

• Bowers Ridge has higher Sm/Yb than the Aleutian Arc, but similar La/Sm. Whereas, Rat

Island falls within the lower part of the Aleutian Sm/Yb and La/Sm ranges, indicating

more residual garnet in the Bowers Ridge source due to deeper melting and or subduction

erosion or slab melting.

2) Rat Island transitioned from the larger volume, tholeiitic volcanism of the Gunners Cove

Formation (NW 2/3 of island) to the smaller volume, calc-alkaline volcanism of the Rat

Formation (SE 1/3 of island) at ~15 Ma, which is also seen on Kiska. The likely cause of this is deeper melting and more water input from subduction due to a shift in tectonics at ~15 Ma

(Minyuk and Stone, 2009), possibly resulting in a change from a transtensional to a transpressional tectonic environment. This is supported by the following observations:

• Ba/La ratios are higher for the Rat Formation samples, indicating more water input from

subduction and the compatible nature of Ba in the amphibole and feldspar of the more

evolved porphyries of the Rat Formation, which are rare in the Aleutian Arc.

• Similar Sr and Nd isotope values indicate that the source was the same and rules out

changes in sediment input.

• Higher La/Sm in the Rat Formation with the same Sm/Yb indicates amphibole

fractionation and smaller degrees of melting for the Rat Formation versus the Gunners

Cove Formation, perhaps reflecting an increase in the water content or thickened crust.

• Lower Sr concentrations, Eu anomalies, and depletions in Hf and Zr are most likely due

to the fractionation of plagioclase and accessory phases such as zircon, respectively.

189

3) The same northward migration of volcanism and transition from tholeiitic to calc-alkaline volcanism observed elsewhere in the Aleutians also occurred in the Rat Island Block. The chemical changes seen through time affect the entire length and width of the arc. Rat Island and the entire Rat Island Block fall within the chemical variations of the Aleutian Arc as their position along the arc would predict.

• The Rat Islands do not form the classic linear array of volcanism seen in the rest of the

arc, but they do seem to exhibit the same northward migration of volcanism seen in the

rest of the Aleutian Arc. The exception to this is Rat Island where the younger Rat

Formation covers the southeastern third of the island and the older Gunners Cove

Formation covers the northwestern two thirds of the island. This northward migration is

most likely the result of subduction erosion slowly eroding the trench. Thus, the volcanic

front is approximately the same distance from the trench through time, but the position of

the trench moves (and thus the volcanic line) due to subduction erosion of the over-

ridding plate.

• There is no trend in chemistry with respect to a volcanic centers east-west placement

within the block or the middle versus the edge of the block. In regards to distance from

the trench, there is no trend in Sr or Nd isotopic ratios, Ba/La, or Ta/La. However, with

increased distance from the trench, there is a slight decrease in Sm/Yb that then increases

again for Segula and Semisopochnoi. A similar but opposite trend is seen for La/Sm and

Th/La, with a slight increase followed by a decrease for Segula and Semisopochnoi. It

should be noted that islands the same distance from the trench do not always have the

same range of chemical values and that most of these islands have few data points and

190

little age constraint, indicating that this correlation may not hold true for a larger data set

and actual age data.

• The chemical changes and the northward migration of volcanism with time were likely

due to changes in the geometry of subduction, which could change the compressive

forces, depth of melting, input of sediments and fluids, and the amount of subduction

erosion. As exhibited by the data examined in this study, these changes are similar, but

not identical, along the arc as there are significant differences along arc in several

controlling factors, e.g., sediment supply, subduction angle and rate, slab dip, etc.

Alternatively, the northward migration of volcanism with time could be due to

subduction erosion moving the trench northward with time. If the volcanic centers were

approximately the same distance from the trench at the time of eruption, this would

explain the lack of a chemical trend with distance from the trench.

4) There are slight differences in the overall chemical signature of the Aleutian Arc from east to west due to variations in the subduction conditions, and Rat Island and the entire Rat Island

Block fall within this array as their position along the arc would predict.

• Sm/Yb and La/Sm, on average, remain nearly constant along the arc.

• Ba/La shows a slight increase to the east, suggesting a larger sedimentary component or

more fluid input in the east, perhaps related to the more orthogonal convergence and

convergence rate to the east as well as to the supply of sediment decreasing westward.

• Th/La and Ta/La are the same on average for most of the arc; while low values of these

ratios on Attu may be due to the more pronounced strike-slip component of the plate

convergence and the more extensional environment in the western Aleutians.

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87 86 • Sr/ Sr is approximately the same across the arc, but is slightly lower for Attu. εNd

decreases to the east. This indicates a slightly more enriched mantle source to the east,

likely due to a change in the geometry of subduction as well as differences in the

sediment/fluid input and possibly subduction erosion input along the arc.

192

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Chapter 3: A Comparison of Sediment Subduction in the Aleutian and Izu-Bonin-Mariana Arcs Through Time

INTRODUCTION Subduction is a key part of the plate tectonic process that continually shapes and reshapes the surface of the Earth. Its importance is evidenced by the active volcanoes of modern island arcs, associated mega-earthquakes, and the virtual absence of oceanic crust older than ~180 Ma

(with the exception of ophiolites). The latter observation suggests that subduction has been occurring for a significant fraction, and perhaps all, of Earth’s history. Yet, all presently observable island arcs are relatively young (with the exception of fossil arcs and ophiolites).

Indeed, Gurnis et al. (2004) note that nearly half of all active subduction zones initiated in the

Cenozoic. This, together with the documented occurrence of paleo-subduction zones (e.g., western North America) and plate reorganizations documented by paleomagnetic studies, suggests that new subduction zones occasionally initiate.

We cannot observe the entire process of subduction zone initiation at present; hence, investigation of zones of subduction initiation and island arc development must be studied through the older rocks of modern arcs, such as the Aleutian Arc and the Izu-Bonin-Mariana

Arcs (IBM). Unfortunately, these older rocks are generally buried beneath younger volcanics. In the IBM Arc, old lavas are fortuitously preserved and exposed both in the forearc and on the

Kyushu-Palu Ridge, which was rifted from the arc by back-arc spreading (Karig, 1971; Karig et al., 1978; Arculus et al., 2015). Because the locus of volcanism has shifted through time and

Pleistocene glacial erosion has exposed the older rocks in some localities, the Aleutians represent another potentially promising area to examine arc initiation (e.g., Yogodzinski et al., 1993).

Comparison with the modern tectonic configuration and the geochemistry of the younger rocks can be used to investigate the initiation and evolution of these arcs. In this study, we compare the

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Aleutian and Izu-Bonin-Mariana Arcs in order to compare arc initiation around the Pacific at approximately 50 Ma and to investigate how these arcs evolved through time.

The IBM arc system is located in the Western Pacific where the Pacific Plate is subducting beneath the Philippine Plate (Figure 3.1). The oldest ages for the IBM are 51-52 Ma

(Ishizuka et al., 2011), which is approximately when at least the central and eastern Aleutian Arc is thought to have initiated (summary in Scholl, 2007; Kay and Kay, 1985, Jicha et al., 2006;

Yogodzinski et al., 1993). Despite these similarities, the IBM Arc differs in several ways from the Aleutian Arc. The IBM has significant back-arc spreading with a full spreading rate of 3-5 cm/yr (e.g, Stern et al., 2003) for the active Mariana Trough; whereas, the Aleutian Arc lacks back-arc spreading at present, except in the Komandorsky Basin, north of the Komandorsky

Islands (e.g., Yogodzinski, 1993). The crust currently being subducted is older in the IBM (Late

Cretaceous-Early Jurassic) (Miller et al., 2005; Stern et al., 2003) than in the Aleutians (Late

Paleocene-Early Cretaceous) (Scholl, 2007 and references therein). The Aleutian Arc’s mid- crustal composition is more mafic than that of some other arcs, including the tonalite composition of the Izu-Bonin-Mariana Arcs (Suyehiro et al., 1996; Calvert et al., 2008;

Takahashi et al., 2007, 2008). The current slab dip in the IBM ranges from shallower (40°) in the north, to steeper (80°) in the south (see review of Stern et al., 2003); while the slab dip in the

Aleutians ranges from 45° in the east, to 60° in the central arc, to 50° in the west (Ruppert et al.,

2007). The current convergence rates for the IBM range from 3 cm/yr in the south to 6 cm/yr in the north (see review of Stern et al., 2003) and vary from oblique to orthogonal. These convergence rates are lower than those of the Aleutian Arc (5.7-7.5 cm/yr) except for the strike slip regime of the western Aleutians where the orthogonal component is very small. Another difference between the two arcs involves the presence of boninites (primitive, high-Mg andesitic

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Figure 3.1: A) Map of the Northwest Pacific Ocean depicting the location and ages of the Izu- Bonin-Marian Arc, the Bering Sea region, and the Aleutian Arc with key islands labeled (Attu, Kiska, Adak, and Umnak). B) Map of the Izu-Bonin-Mariana Arc and surrounding area. Maps produced using GeoMapApp (www.geomapapp.org).

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Figure 3.1 (Continued):

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lavas with distinctive rare earth patterns) in the IBM that have not been found in the Aleutian

Arc and the presence of adakites (intermediate to silicic, high Sr/Y and La/Yb) principally in the western Aleutian Arc that are not found in the IBM, indicating a difference in magmatic source, processes, and/or tectonics that result in a difference in the geochemistry of the eruption products.

Initiation of the Aleutian Arc (discussed in Chapter 1) differed from that of the Izu-

Bonin-Mariana Arcs. The model presented by Stern et al. (2003) and refined by Stern and Gerya

(2017) has IBM initiation occurring at a transform boundary in an extensional seafloor spreading regime at approximately 50 Ma (Bloomer et al., 1995; Cosca et al., 1998; Stern et al., 2003).

Stern and Bloomer (1992) proposed a N-S trending fracture zone with an E-W spreading regime.

The older denser Western Pacific crust foundered and sunk until asthenosphere could flow over it, causing large scale decompression melting and the eruption of forearc basalts (FABs)

(depleted tholeiites), while the younger crust of the Philippine Plate remained stable and became the upper plate in the developing subduction zone. Ishizuka et al. (2018) add that the presence of remnant arc terrranes on the younger plate also led to differences in the density of the two plates, favoring subduction initiation. As the Pacific Plate continued to sink, it eventually released fluids, which lead to melting of the depleted mantle and the eruption of boninites, followed by the later eruption of low-K rhyodacites (Hickey and Frey, 1982; Stern et al., 1991; Taylor et al.,

1994; Stern et al., 2003; Stern and Gerya, 2017). Subsidence gave way to true subduction at ~43

Ma, magmatism transitioned from the initial broad outpouring to more localized individual volcanic centers, and the forearc lithosphere and sub-forearc mantle stabilized and cooled.

Hickey-Vargas et al. (2018) found that during the first ~3 Ma of subduction, FAB-like basalts were erupted in the back arc due to decompression melting, while FABs and boninites were

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erupted in the forearc largely due to hydrous flux melting. This change to true subduction was

thought to correspond to the change in the direction of Pacific Plate motion from northerly to

westerly at ~43-50 Ma.

Arculus et al. (2015) agree with the assessment that IBM subduction initiation was

spontaneous, but they do not believe that it started at a pre-existing fracture zone. They cite as

evidence the existence of extensive basaltic sheet lavas in the Amami Sankaku Basin, west of the

Bonin Ridge (drill site U1438), with geochemical characteristics indicative of a depleted mantle

source that required fluid fluxing to melt and that have trace element characteristics indicative of

subduction-related magmatism in the early stages of arc development. They conclude that the

IBM subduction initiation at 52 Ma triggered rifting and seafloor spreading on the overriding

plate, which created extensive basaltic arc crust along- and across-strike. In addition, there is no evidence of the uplift associated with the compression inherent in the induced subduction process, but there is evidence of spreading and rifting, which would occur in a spontaneous subduction situation that lacks compression (Arculus et al., 2015). This spontaneous subduction initiation is thought to have been triggered by subsidence of the older denser Pacific Plate in relation to the younger more buoyant Philippine Sea Plate along a system of transform faults or fracture zones between the two plates. The arc volcanoes later developed on this early crust as volcanism became more localized near the trench. Based on the tectonic and geochemical constraints of the area, they conclude that the spontaneous subduction initiation of the IBM did not occur at a preexisting fracture zone.

At approximately 30 Ma, the arc rifted and the Parece Vela Basin formed (Taylor, 1992;

Stern et al., 2003). In the north, rifting started at approximately 25 Ma and propagated southward, resulting in the formation of the Shikoku Basin (Kobayashi et al., 1995; Stern et al.,

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2003). The two spreading systems (Parace Vela Basin and Shikoku Basin) met at ~20 Ma and

continued spreading until ~15 Ma. During this rifting stage, magmatism waned along the arc and

did not resume until the later stages at ~20-17 Ma when back-arc seafloor spreading had replaced rifting (Lee et al., 1995; Taylor, 1992; Stern et al., 2003). At ~15 Ma, the northernmost part of the arc underwent collision with Honshu, which Stern et al. (2003) propose was the result of new subduction along the Nankai Trough.

The last stage of Mariana back-arc evolution was the rifting of the Mariana Trough after

7 Ma (Karig, 1971; Karig et al., 1978; Arculus et al., 2015) and back-arc seafloor spreading from

3-4 Ma to present day (Bibbe et al., 1980; Yamazaki and Stern, 1997; Stern et al., 2003). The removal of the old arc by back-arc spreading results in the age of the present Mariana Arc volcanoes being 3-4 Ma or younger, whereas the Izu segment volcanoes may be as old as 25 Ma.

Rifting in the Izu segment began at ~2 Ma and continues to the present day (Taylor, 1992; Stern et al., 2003).

Following the eruption of the FABs, boninites were erupted, which are rare in other island arcs. Boninites are defined as andesites (SiO2 > 53 wt.%) with high MgO (> 5 wt.%), Cr

(500 ppm), and Ni concentrations; low TiO2 (< 0.6) and very low incompatible element

concentrations; contain orthopyroxene phenocrysts or microphenocrysts, and lack feldspar

(Crawford et al., 1981; Taylor et al., 1994; Arndt, 2003). Given that the boninites are

stratigraphically located immediately above the arc tholeiites, the eruption of these boninites is

thought to be the result of melting of a peridotite mantle source that was depleted by relatively

high degrees of previous melting and ridge volcanism (Crawford et al., 1981). Arndt (2003)

observed that the high degrees of shallow melting required to produce boninites could only be possible in the presence of large amounts of subducted slab-derived hydrous fluids, which is

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supported by the high Al2O3/TiO2, low CaO/Al2O3, and low FeO of boninites. In order to obtain the characteristic U-shaped REE profile of LREE enrichment superimposed on a depleted signature, an enriched fluid component must be added to the mantle wedge (Taylor et al., 1994;

Arndt, 2003). Crawford et al. (1981) note that boninites seem to be associated with the early stages of the magmatism that split the Palau-Kyushu Arc and formed the Parece Vela Basin.

They speculate that boninites may only occur in the early stages of back-arc rifting; however, the stratigraphic relationship between the boninites and tholeiitic basalts of such episodes is unclear.

Stern et al. (1991) investigated possibilities for the fluid component that enriched the depleted harzburgite mantle. Based on isotopic modeling, they concluded that sediments could not be a significant source of enrichment and that the fluid component was the result of dehydration of oceanic crust. Bloomer and Hawkins (1987) concluded that the fact that boninites only appear in the initial stages of IBM subduction indicates that the conditions of this early subduction are significant to the generation of boninites. They believe that the presence of the young hot West

Philippine Basin crust and the release of volatiles at a shallow depth into a shallow reservoir of depleted mantle are key factors in the production of boninites. Given the shallow isotherms of the young subduction zone, depletion of the mantle source by early volcanism would have occurred at fairly shallow depths, and the slab would have also been dehydrated at shallow depths within this depleted reservoir. With time, the isotherms migrated to deeper depths, which would have been accompanied by melting and dehydration also moving to greater depths and into the less depleted mantle (Bloomer and Hawkins, 1987).

The tectonic history of the two arcs (see Chapter 1 for a summary of the history of the

Aleutian Arc) is significantly different as is the geochemistry of their products. For example, there have been no boninites found in the Aleutians despite its having a similar initial phase of

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rapid growth of the arc massif (Kay and Kay, 1994). In this study, we compare the Aleutian and

Izu-Bonin-Mariana Arcs to investigate the processes that initiated subduction around the Pacific at approximately 50 Ma and to investigate how these processes progressed through time to form the various geochemical signatures of these arcs.

GEOCHEMISTRY OF THE ALEUTIAN AND IZU-BONIN-MARIANA ARCS

We can investigate differences in subduction zone initiation and evolution around the

Pacific at ~50 Ma by comparing the geochemistry of the IBM and Aleutian Arcs through time.

Subduction initiated at both arcs at about this time, yet they have distinctly different geochemical and tectonic histories. In order to represent the full range of Aleutian geochemistry in an orderly fashion, we take several well studied islands as being representative of the different areas of the arc. We take Attu as being representative of the western Aleutians, Kiska as an intermediate central island, and Adak and Umnak as representative being of the central and eastern intra- oceanic portions of the Aleutian Arc. These islands were chosen because of their locations along the arc and the fact that there is a reasonable amount of data available for these islands compared to others islands. In both arcs, sampling of older volcanics and plutonics (due to the scarcity of eruptive events at 10-25 Ma; Taylor, 1992; Bloomer et al., 1995) is too sparse to reconstruct the continuous evolution of the IBM and Aleutian Arcs through time, particularly from 10-25 Ma.

Therefore, we consider the data in two broad age groups. We use >30 Ma samples to investigate the arc initiation and <6 Ma samples to investigate the arcs in their present states. Data for intermediate ages is too sparse to draw reasonable inferences. In order to focus on subduction zone factors (e.g., source and initial melt characteristics vs. fractional crystallization, mixing, and assimilation during eruption), we have limited our data sets to SiO2 < 60%.

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Figure 3.2: SiO2 vs. K2O discrimination diagram. Black lines delineate low, medium, and high- potassium groupings (after Gill, 1981). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. References are as follows: Attu: (includes data from Yogodzinski et al., 1993; Shelton, 1986; and Rubenstone, 1984); Kiska (includes data from George et al., 2003); Adak (Coats, 1952; Conrad et al., 1983; Debari et al., 1987; Kay and Kay, 1985; Kay and Kay, 1994; Kay et al., 1982; Marsh, 1976; Myers et al., 1985; Walker, 1974; Kay et al., 1986; Yogodzinski et al., 2010; Coats, 1952; Debari et al., 1987; Kay and Kay, 1988; Kay and Kay, 1985; Kay and Kay, 1994; Marsh, 1976; Myers et al., 1985; Von Drach et al., 1986); Umnak (Singer et al., 2007); IBM (EarthChem and GEOROC databases).

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Figure 3.3: Th/La vs. K2O diagram. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. Data references as in Figure 3.2.

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Figure 3.4: Chondrite normalized spider diagrams. Normalization values from Sun and McDonough (1995). A) Izu (green), Bonin (black), and Mariana (purple). B) Adak (orange), Umnak (gray), Attu (red), and Kiska (pink). Data references as in Figure 3.2.

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The Izu-Bonin-Mariana and Aleutian Arcs have both erupted tholeiitic and calc-alkaline magmas. The Aleutian Arc has higher on average K2O than the IBM (Figure 3.2). The Aleutian samples are predominantly medium-K with a few (mostly altered) samples ranging into the high and low-K fields. The Izu segment of the IBM is low-K except at high silica; and the Mariana and Bonin segments range from low-K to medium-K, with a few high-K samples in the modern

Mariana segment. Based on the classification described by Sun and Stern (2001), the high-K samples in the modern Marianas are not shoshonites. A few lower silica samples in the Bonin and Mariana segments range to the high-K field, but given their older ages (>30 Ma), the high

K2O of these samples may be due to alteration, as is also seen in the Sr isotopes of the Bonin segment.

The K2O of Adak, Attu, and the Modern Mariana and Izu segments does not vary with

Th/La. The K2O of Kiska increases with increasing Th/La; however, it should be noted that this trend is only based on four data points (Figure 3.3A). The K2O of the ancient Mariana and Bonin segments and Attu does not vary with Th/La. While the K2O of the >30 Ma Kiska and Adak samples vary only slightly with Th/La (Figure 3.3B).

Rare earth element (REE) patterns in lavas from the Mariana segment have similar light rare earth element (LREE) enrichments to Adak, but are not as steep (Figure 3.4). Attu and Kiska look more like the Izu segment and the upper half of the concentration range of the Bonin segment. The exception to this is the younger ≤ 5 Ma Mariana calc-alkaline samples, which have much steeper more Adak-like patterns and lower heavy rare earth element (HREE) contents. The

Late-Dikes (16 Ma) of Attu have flat to slightly enriched patterns like the Mariana segment, but have higher concentrations overall and do not match any other Aleutian or IBM REE patterns.

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The older samples (>30 Ma) from the Mariana-Bonin and Aleutian Arcs define distinct fields in (La/Sm)N-(Sm/Yb)N space that overlap in their most depleted endmembers, which fall in the more depleted half of the North Pacific MORB field (defined by data downloaded from

PetDB www.petdb.org). At >30 Ma, the Mariana segment (La/Sm)N ratios range from less than

1, similar to North Pacific MORB, to 3.5, similar to those of the Adak and Umnak samples. The older (>30 Ma) samples from the Bonin segment have more restricted and lower on average values (LREE depleted) compared to those of the Mariana segment. The (La/Sm)N and (Sm/Yb)N values of the central Aleutians are higher than those of the western Aleutians at >30 Ma, with the lowest values in the western Aleutian samples plotting within the North Pacific MORB field

(Fig. 3.5). Overall, for all of the age groups, the Aleutian samples are more LREE enriched and exhibit more fractionated patterns than the IBM samples do.

For the most recent volcanics (<5 Ma), the Mariana segment’s (La/Sm)N and (Sm/Yb)N ratios stretch from the North Pacific MORB field through the Aleutian range, with a trend of increasing (La/Sm)N with increasing (Sm/Yb)N that is much more distinct than the trend for the older Mariana samples. Adak and Umnak have higher on average (La/Sm)N and (Sm/Yb)N values than those of the Western Aleutians (except for the adakites and dredge samples of submarine volcanism, which are not considered in this study; Yogodozinski et al., 2015;

Yogodzinski et al., 2017) and the IBM, especially the Izu segment. At <5 Ma, the Western

Aleutian samples have higher on average (La/Sm)N and (Sm/Yb)N than the IBM samples (with the Mariana field having a clump of more enriched samples) and fall into the lower end of the

Adak field. While the lower Umnak values overlap with the western Aleutian samples and the higher Izu values.

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Figure 3.5: Sm/Yb vs. La/Sm normalized to C1 chondrite (McDonough and Sun, 1995). North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. Data references as in Figure 3.2.

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Figure 3.6: Th/La vs. Ba/La diagram. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. Data references as in Figure 3.2.

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The Ba/La ratios of the young (<5 Ma) lavas from the Izu and Mariana segments overlap

the MORB range, but extend to much larger values as well (Figure 3.6). The younger Aleutian

samples do not overlap the MORB range but generally do not extend to ratios as high as those of

the Izu and Mariana lavas. The Ba/La values of the Mariana and Izu segments overlap each

other, with the Izu segment values being slightly higher on average than those of the Mariana

segment. On average, the Ba/La values of the Aleutian and IBM arcs are the same, but the IBM

values extend over a greater range of values (to higher and lower values). The modern IBM

samples have lower Th/La than the Aleutian samples, and once again, the Izu and Mariana

segments have a MORB-like endmember that the Aleutians lack. The Kiska and Attu samples

are intermediate between the IBM data and that of Adak and Umnak.

Th > 30 Ma Mariana samples plot slightly higher on average in Ba/La and Th/La than the

Bonin samples, though both have a MORB-like endmember. The Ba/La of the older samples

from both arcs are roughly similar; however, the older Attu and Kiska samples exhibit a much

larger range of values, perhaps in part due to secondary alteration. The Th/La ratios of the IBM

(Figure 3.6) are lower on average than those of Adak and Umnak, but they share a considerable

amount of overlap. The older Attu samples also have a Ba/La-Th/La depleted endmember

(MORB-like) (30-35 Ma Plutons). The older Kiska samples have lower Th/La, similar to Attu, but do not extend into the MORB field in Ba/La.

The Izu and Bonin segments have exclusively arc-like Ta/La ratios, while the Mariana samples have a few MORB-like samples, which represent the early forearc tholeiites and upwelling from back-arc spreading of the Mariana Trough (Mariana Trough samples were not considered in this study). In contrast, the older Aleutian Arc samples span the range from an arc- like to a MORB-like setting on the ridge itself, although most prominently on Attu where the

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. ), and Mariana Mariana and ), (purple). Data ack Izu (green), Bonin (green), Bonin Izu (bl (defined by data downloaded from PetDB www.petdb.org) data PetDB downloaded from by (defined

blue diagram. North Pacific MORB field in in field MORB Pacific North diagram.

Nd ε Sr vs. 86 Sr/ 87 : 7 references as in Figure 3.2. Figure 3. (pink), Adak AttuKiska (red), and field (orange), Umnak and and (grey); 220

subduction has a prominent strike-slip component and there is evidence of incipient intra-arc

rifting that never progressed to the back-arc spreading stage (see discussion in Chapter 1). The

modern Aleutian and IBM samples have similar Ta/La, and the modern Aleutian samples are

exclusively arc-like.

Figure 3.7 compares the Nd and Sr isotopic compositions of the two arc systems (because

of the paucity of data, 5-15 Ma samples from Attu and Kiska are included in this plot). The high

87Sr/86Sr ratios of many of the older Attu and Kiska samples and perhaps some of the IBM

samples reflect secondary hydrothermal seawater alteration (e.g., Stern et al., 2003; Reagan et

al., 2010). Emphasis is thus placed on the εNd values, which are robust with respect to secondary

processes. The Aleutian samples do not separate into different trends or areas with age.

87 86 The εNd and Sr/ Sr values of the older samples from the Mariana segment overlap the lower part of the MORB range and extend to much more enriched signatures (lower εNd and higher 87Sr/86Sr) than the Aleutians. The modern Mariana samples overlap with the less enriched

older Mariana samples and overlap the Aleutian field, but they extend to more enriched εNd. The

87 86 Bonin samples with the highest εNd and lowest Sr/ Sr barely overlap the Aleutian range and

87 86 extend to lower εNd, although not quite as low as the Marianas, and particularly high Sr/ Sr

values. In contrast, the Izu samples show a much more restricted range of isotopic compositions,

87 86 overlapping the unaltered part of the Aleutian range, with slightly lower in εNd and Sr/ Sr on

average than the Aleutian samples. The fact that the IBM samples span such a wide range of

87 86 Sr/ Sr values (especially in the older samples), but span a much smaller range of εNd values,

indicates that the older samples have undergone seawater alteration, and like the older Aleutian

samples, many, and perhaps most, of their 87Sr/86Sr is no longer primary. The values that

correlate with εNd are likely reliable, while the values that scatter away from this correlation are

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likely altered. In summary, the IBM isotope ratios overlap those of the Aleutian samples, but

extend to more enriched signatures than those of the Aleutian Arc.

The Mariana and Bonin segments exhibit a trend of increasing Ba/La with decreasing εNd,

while the Izu segment and the Aleutian samples exhibit increasing Ba/La at a relatively stable εNd

(Figure 3.8). The IBM samples form a vague, very scattered trend of increasing (La/Sm)N with

decreasing εNd (Figure 3.9). The Mariana samples just overlap the εNd of the MORB field and

mostly encompass the Izu and Bonin fields. On average, the modern Mariana samples are more

LREE enriched than the other IBM samples, are more enriched than the western Aleutian

samples, and completely overlap with the eastern/central Aleutian samples. Though it is slightly vaguer and just as scattered, the inverse (La/Sm)N verses εNd correlation is also seen for the

Aleutian samples. However, the Aleutian trend does not extend to such low εNd values as the

IBM trend. In the Aleutian Arc, there is a trend of increasing Th/La with decreasing εNd (Figure

3.10). Whereas, in the IBM Arc, Th/La is constant with decreasing εNd. The range in εNd is quite

a bit larger for the Mariana segment than for the Bonin and Izu segments, with the Izu values

being within the Aleutian range. Once again, on average, the modern Mariana samples have

lower εNd (more enriched) than the older samples.

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Figure 3.8: Ba/La vs. εNd. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. Data references as in Figure 3.2.

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Figure 3.9 La/Sm normalized to C1 chondrite (McDonough and Sun, 1995) vs. εNd. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. Data references as in Figure 3.2.

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Figure 3.10: Th/La vs. εNd. North Pacific MORB field in blue (defined by data downloaded from PetDB www.petdb.org). Kiska (pink), and Attu (red), Adak field (orange), and Umnak (grey); and Izu (green), Bonin (black), and Mariana (purple) fields. Data references as in Figure 3.2.

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COMPARISON OF THE ALEUTIAN AND IZU-BONIN-MARIANA ARCS

Differences in composition between the IBM and Aleutian Arcs could reflect the

following factors:

• a difference in the amount and/or composition of the sediment being subducted beneath

the arcs;

• a difference in the amount of slab-derived aqueous fluid supplied to the mantle wedge

beneath the arcs;

• a difference in the crustal and lithospheric thicknesses of the arcs, which would affect the

height of the melting column (Plank and Langmuir, 1988) leading to

o lesser extents of melting of the mantle wedge beneath thicker arcs

o and greater fractional crystallization within thicker arcs

• and a difference in the mantle source under the two arcs, which was the same before

subduction began.

In the following discussion, these possibilities will be considered separately.

Sediment subduction beneath the arcs at >30 Ma: Although some studies suggest Indian

Ocean Mantle as the mantle source of the IBM (Yogodzinski et al., 2018), many others have used NMORB and Pacific MORB (e.g., Woodhead, 1989; Elliot et al., 1997; Peate and Pearce,

1998; Stern et al., 2003; Pearce et al., 2005). Assuming for the moment that the mantle source under the two arcs was the same before subduction began (Pacific MORB), then any differences in source characteristics must be due to the subduction process, principally to sediment subduction. Differences between source characteristics, i.e., sediment addition, are best constrained by isotope ratios. Both arcs exhibit variably elevated 87Sr/86Sr relative to MORB.

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Excluding secondary alteration effects, this can, and no doubt at least in part does, reflect the

contribution of seawater-altered oceanic crust. Either variably altered crust or variable amounts

87 86 of this material would produce variable Sr/ Sr. In contrast low εNd values relative to MORB

can only be due to contributions of subducted sediment.

Both the amount and composition of the sediment subducted will influence the Sr-Nd

isotope systematics (Figure 3.7). Though the sediments currently being subducted provide us

with information on only the more recent volcanism, they can still be useful in considering the

differences between the two arcs. The sediments currently subducting beneath the Aleutians have

less enriched isotopic signatures, have lower Nd concentrations, are more enriched in Sr, Ba, and

K2O, and have higher Th/La and Ba/La ratios than the sediments subducting beneath the Izu-

Bonin and Mariana arcs; but their La/Sm and Sm/Yb ratios are fairly similar (see summary in

Plank, 2014; Schaen et al., 2016). The Aleutian, Izu-Bonin, and Mariana sediments have the

87 86 following respective values: εNd of -0.158, -5.8, and -2.3; Sr/ Sr of 0.70635, 0.70956, and

0.70617; Nd of 19.1, 25.2, and 21; Sr of 245, 136, and 161; K2O of 2.11, 1.16, and 1.36; Th/La

of 0.305, 0.188, and 0.126; Ba/La of 115, 37.78, and 14.95; (La/Sm)N of 4.42, 4.78, and 4.42;

and (Sm/Yb)N of 1.24, 1.29, and 1.53 (see summary in Plank, 2014; Schaen et al., 2016). The

trace elements of the Mariana, Izu-Bonin, and Aleutian sediments are roughly similar, but their

concentrations of several key elements vary. These elements include Ba, Th, K, Sr, and Pb. The

Izu-Bonin sediments are more enriched in REEs than the Aleutian and Mariana sediments, while

the Aleutian sediments are more enriched in Ba, Th, and K. The Izu-Bonin sediments have the

largest Nb-Ta and Zr-Hf depletions. In general, the enrichment of the subducting sediments from

most to least enriched is as follows: LREE - Izu-Bonin, Mariana, Aleutians; HREE – Izu-Bonin,

Aleutians, Mariana; HFSE – Mariana/Aleutians, Izu-Bonin; and LILE – Aleutians, Izu-Bonin,

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Figure 3.11: Primitive mantle normalized spider diagrams of Mariana (purple), Izu-Bonin (green), and Aleutian (red) sediments. Normalization values from McDonough and Sun (1995). Sediment data from Plank, 2014.

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Mariana (see summary in Plank, 2014; Schaen et al., 2016). Figure 3.11 displays these values on

a primitive mantle normalized spider diagram.

The enrichment of the Aleutian sediments in Th relative to La compared to those of the

IBM system (Figure 3.11) likely reflects the more terrigenous nature of the sediments in the

Aleutians, which in turn reflects the proximity of a continental source (Alaska) and the effects of

Pleistocene glaciation (Kay and Kay, 1994; Schaen et al., 2016; Kay et al., 2018). The Aleutian

sediment is also richer is Ba and K2O, although not necessarily due to continental sediment input

(authigenic barite can be an important component of some sediments) (Torres et al., 1996; Suess

et al., 1998).

These differences in sediment composition may in turn explain some of the

compositional differences between the arcs. Increasing Ba/La with decreasing εNd in the Bonin

Arc (Figure 3.10) indicates that the Ba/La is controlled by sediment-derived fluid input, while a lack of correlation, as is the case in the other arcs, indicates that the Ba/La is controlled by ocean crust-derived fluid input (see correlation with 10Be and Ba in Kay and Kay, 1994), assuming it is

not at least partially caused by later seawater alteration.

As a consequence of the high Nd concentration and low εNd of the sediment being

subducted beneath the Bonin Arc, the isotopic compositions of the magma in this arc are

particularly sensitive to the sediment component of their source (Fig. 3.7). Three to five percent

sediment would explain the large range of εNd for the Bonin Arc magmas. In contrast, due to the

higher εNd and lower Nd concentration of the sediment being subducted beneath the Aleutians,

relatively more sediment would be required in the Aleutians to produce the same isotopic

compositions as seen in the IBM arcs. A one or two percent sediment contribution would explain

the most extreme Aleutian compositions, and the sediment contribution appears to be much less

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than that for most magmas, which is consistent with the finding of other studies (e.g., Kay, 1980;

George et al., 2003).

The εNd values for the older Aleutian samples extend barely beyond the MORB field,

while their 87Sr/86Sr ratios are systematically higher than MORB. These elevated 87Sr/86Sr ratios may also in part be due to a contribution of subducted seawater-altered oceanic crust. However, extensive seawater interaction has affected many of the older samples in the western Aleutians, as evidenced by greenschist metamorphism (Rubenstone, 1984; Yogodzinski et al., 1993) as well as their Sr isotope compositions being too high to be explained by even a 100% AOC endmember. Based on the mixing curves in Figure 3.7, the εNd of the Aleutian samples can be

explained by a 97%DM and 3%AOC endmember and the addition of <1% volcanoclastic

sediment (i.e., sediment derived from arcs and seamounts).

The Bonin segment has similarly elevated Sr values for some samples that cannot be

produced by the addition of AOC or sediment. Thus, we conclude that these high Sr isotope

values are also the result of later seawater alteration, but those that correlate with εNd (the lower values that fit the mixing line) are relatively unaltered. These samples can be modeled with an endmember of 97%DM and 3%AOC with the addition of 1% Izu-Bonin segment sediment, while up to 3% addition of sediment can account for the entire range of εNd (overlooking the

elevated Sr due to alteration). In contrast, the Mariana Sr isotopes can be modeled by mixing

between a DM endmember and 10% Mariana sediment alone, with little apparent need for

seawater-altered oceanic crust. This amount of sediment mixing is consistent with the results of

previous studies (e.g., Tera et al., 1986; Ryan and Langmuir, 1988; Woodhead, 1989; Elliot et

al., 1997; Peate and Pearce, 1998; Pearce et al., 2005; Wade et al., 2005; Woodhead et al., 2005;

Chauvel et al., 2009).

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It should be noted that the boninites have low εNd values despite the sediment

contribution being modest, <3%, which reflects the very low εNd (-5.8) (Plank, 2014) of the

sediments being subducted. In addition, the particularly depleted boninite mantle source

(Crawford et al., 1981) would need less sediment to dramatically shift isotopic compositions.

Another point of importance is that the boninitic samples were produced by low pressure melting

of a depleted mantle source with fluid fluxing, and thus, the more depleted source would require

less contribution of sediment and fluids to enrich it to the values observed (Crawford et al., 1981;

Stern et al., 1991; Taylor et al., 1994; Arndt, 2003). For example, if the Bonin mantle were 50%

more depleted in Nd and Sr, only 0.5-2% sediment addition would be required to explain the data. If the Bonin mantle were 10% more depleted in Nd and Sr, only 0.66-3% sediment addition would be required to explain the data. This could also partially account for the fact that the Bonin samples have a similar range of εNd and Ba/La values (Bonin Sr values are higher but altered),

but the mixing lines indicate a much smaller percent of sediment input (3.25% Bonin vs. 10%

Mariana). However, how much of this effect is due to the differences in the composition of the

sediments and how much is due to the differences in the depletion of the source is not clear, due

to the fact that the samples represent a span of ages, the fact that the mantle would have been

depleted over time rather than at one instant in time, and the fact that the time at which the

mantle depletion occurred is not well constrained.

It is significant that the mixing lines for the older IBM samples (produced by a

subduction-initiation ridge; Pearce, 2014) do not indicate the addition of volcanoclastic

sediment, but the younger Mariana samples do, indicating that volcanoclastic sediment was not

likely present (at least in large quantities) on the subducting plate in the past (Stern et al., 2003;

Elliot et al., 1997). Thus, subducted sediment had a major, though variable, influence on ancient

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IBM geochemistry, with a more enriched sediment supply entering the early arc system. In

contrast, sediment played at best a small role in the early Aleutian Arc, with the geochemical

signature being primarily the result of oceanic crust-derived fluids, which is consistent with the

findings of Yogodzinski et al. (1993).

The older IBM samples also exhibit more enriched isotopic signatures than the modern

IBM samples, particularly in the Mariana segment. This change in isotopic composition in the

IBM system likely reflects a change in the nature of the sediment being subducted beneath the

IBM from terrigenous- and pelagic-dominated to a sedimentary column dominated by

volcanoclastic material derived from the arc itself, as the latter has less enriched isotopic

signatures than the other sediments (Peate and Pearce, 1998; Ishikawa and Tera, 1999; Stern et

al., 2003) and the modern Mariana samples have higher εNd. This would be particularly relevant

for the Mariana segment as there is more volcanoclastic sediment present on the subducting plate

in the south where the most voluminous volcanism occurred (Stern et al., 2003; Reagan et al.,

2010). This hypothesis is supported by the mixing lines, which indicate volcanoclastic sediment

involvement in the modern Mariana and Izu segments, but not in the older Bonin and Mariana

segments.

In contrast, the Aleutian isotope ratios reflect a much smaller sediment input and exhibit

a relatively smaller change with time (Kay et al., 1990; Kay and Kay, 1994; Schaen et al., 2016)

compared to the IBM. The older Aleutian samples have high εNd and are incompatible element

depleted, which is indicative of less sediment involvement than is seen in the IBM. We have no samples of the sediment that was subducted at >30 Ma; fitting mixing lines based on modern sediments allows some estimate of how much and what kind of sediment was subducting at that time. In particular, based on the mixing lines that fit the data best, the Adak samples have more

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volcanoclastic involvement (97%DM-3%AOC + <1% volcanogenic sediment) and Attu has less

sediment involvement and the sediment is not volcanogenic (97%DM-3%AOC + <0.5%

Aleutian sediment). This is more a reflection of changes in the subduction geometry and volcanic

output along arc than changes with time. The convergence is more perpendicular and more

volcanism occurs to the east, which results in more volcanogenic sediment being produced and

subducted, due to erupted tephra, ash, scoria, etc., and erosion of the volcanic material. Thus, the

IBM experiences changes in chemistry due to changes in subduction input through time and

along arc, while the Aleutian Arc is primarily controlled by the more oblique subduction and

lower sediment supply to the west.

Subducted sediment-derived fluids and or/melts added to the mantle wedge would

influence not only the isotopic composition, but also the incompatible element abundance. Thus,

from the foregoing discussion, one would infer that the mantle sources of the IBM magmas,

particularly those of the older Mariana and Bonin arcs, were more incompatible element-

enriched than the mantle source of both the ancient and modern Aleutian Arc. However,

assessing the sediment contribution to elemental abundances is more difficult because solid- silicate liquid and solid-aqueous fluid partitioning will result in fractionations of elemental ratios such as Ba/La and La/Sm. In contrast, assuming a mantle wedge composition that is uniformly similar to the MORB source, Nd isotope ratios are a comparatively simple function of the amount of sediment added. The extent to which incompatible elements are a function of sediment addition can consequently be examined by assessing the extent to which incompatible elements and their ratios correlate with εNd.

This can be evaluated by determining the correlation coefficients, r, and whether they are

statistically significant for the available number of data points. The square of the correlation

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coefficient (r2) indicates the fraction of the variation in the y variable, e.g., La/Sm that is related to the variation in the x variable, e.g., εNd. Since it is assumed that the variation in εNd is due to sediment addition, the r2 value provides an estimate of how much of the variation in the incompatible element ratios is due to sediment addition. These effects are explored below.

Table 3.1: Correlation coefficients for various trace element ratios versus εNd

La/Sm Th/La Ba/La K2O r n r n r n r n > 6 Ma Attu ------> 6 Ma Kiska ------> 6 Ma Adak ------< 5 Ma Izu 0.5766 63 - - 0.3445 63 - - < 5 Ma Mariana 0.4809 65 0.5449 64 - - 0.7898 65 > 30 Ma Mariana 0.7746 66 - - 0.3820 66 0.3860 64 > 30 Ma Bonin - - - - 0.5940 24 0.8447 31

Table 3.1 lists the correlation coefficients for the cases where the correlation is statistically significant at the 1% level between εNd versus (La/Sm)N, Th/La, Ba/La, and K2O as well as the number of data points used to determine the correlation (correlation coefficients were not determined for the <6 Ma Attu, Kiska, and Adak samples or the Umnak samples due to a lack of sufficient data points). The existence of a statistically significant correlation between εNd and (La/Sm)N, Th/La, Ba/La, and K2O (Figures 3.9, 3.10, 3.8, 3.2, respectively) suggests that sediment addition is a controlling factor in some cases. In other cases, lack of a significant correlation suggests that sediment addition is primarily controlled by some other factor or process.

As with the <5 Ma samples, we can use correlations between εNd and various ratios to determine the extent to which sediment addition effected the incompatible trace elements. Based on the r2 values, sediment melt addition is responsible for 60%, 14.6%, and 14.9% of the

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Mariana segment’s variation in (La/Sm)N, Ba/La, and K2O, respectively (Figures 3.9, 3.8, 3.2,

respectively), although the Th/La correlation indicates that sediment melt is not a significant

factor (Figure 3.10). Sediment addition apparently effects the Bonin segment’s Ba/La and K2O

more (35.3% and 72.4%, respectively), but it has no significant effect on the (La/Sm)N or Th/La.

In the Aleutians, sediment addition apparently does not have a significant effect on the

incompatible trace elements.

Overall, sediment is much less important in explaining incompatible trace element

variations in the Aleutians than in the Mariana and Bonin segments, which is consistent with

their respective sediment supplies (Mariana = 10%; Bonin = 4%; Aleutians <1%; Figure 3.7).

Other factors affecting the incompatible trace element enrichments of these areas will be

discussed in subsequent sections.

Sediment subduction beneath the arcs from 0-5 Ma: The modern Mariana samples exhibit

87 86 a distinctly larger range of εNd (and Sr/ Sr) than the modern Aleutian samples, while the Izu

87 86 samples exhibit a similar range of εNd, but have slightly higher on average Sr/ Sr then the

modern Aleutian samples, with only a slightly broader range in 87Sr/86Sr. This implies

considerably more sediment subduction in the Mariana segment than in the Izu segment and the

Aleutians.

In order to determine the type and amount of sediment added to each arc through time,

87 86 Sr/ Sr-εNd sediment mixing lines were created using four different sediments (Marina, Izu-

Bonin, Aleutian, and volcanoclastic sediment) (Plank, 2014; Woodhead, 1989). The mixing lines

that fit the data best are shown in Figure 3.7. A depleted mantle endmember with Nd of 0.713,

87 86 εNd of 9.25, Sr of 9.8, and Sr/ Sr of 0.7027 (Salter and Stracke, 2014); an altered oceanic crust

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87 86 endmember with Nd of 9.22, εNd of 9.25, Sr of 110, and Sr/ Sr of 0.704 (White and Klein,

2014); and endmembers made from various mixing ratios of these two endmembers were used as

the mantle endmember of the sediment mixing lines. The composition of the volcanogenic

87 86 sediment used is as follows: Nd of 44.22, εNd of 2.302, Sr of 176.21, and Sr/ Sr of 0.7045. The values for the other sediments are listed above (Woodhead, 1989).

As can be seen from the mixing curves in Figure 3.7, the <5 Ma data has a considerable spread in data, which makes it hard to fit a single mixing line. This data spread could signify variation in the sediment composition either over time or along the arc segment. The lack of a trend in 87Sr/86Sr with Ba/La indicates that the younger samples do not experience scatter due to

later seawater alteration. If the scatter were caused primarily by variations in the sediment

composition, we would expect to see a more wedge shaped field that converges on a common

endmember and spreads out between two mixing lines, which is what we see in the young

Mariana data. If the scatter were caused by variable addition of AOC in the endmember, we

would expect to see a wider scatter pattern that does not converge at a common endmember, but

spreads out along the DM-AOC mixing line, which is what we see for the Izu segment. The

Adak and Umnak data can be modeled using an endmember that is 97% depleted MORB (DM)

and 3% altered oceanic crust (AOC) and adding <1% volcanoclastic sediment, which is

consistent with the findings of other studies (e.g., Kay, 1980; George et al., 2003). Though there

is little data for the Kiska and Attu samples (excluding submarine samples), the Kiska samples

appear to roughly fall along the mixing line with a DM endmember and addition of 2% Aleutian

sediment, while the Attu samples roughly fall along the mixing line with a 97%DM and 3%AOC

endmember with <1% Aleutian sediment added. The Izu segment was arguably the hardest to fit

due to the scattering of the data, but the best fit appears to be an endmember of 93%DM and

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7%AOC with 1-1.5% volcanoclastic sediment added. The modern Mariana segment data can be modeled using an endmember of 97%DM and 3%AOC, but due to the wide Sr isotopic variation, the data is bracketed by the lines for Mariana sediment (5%) and volcanoclastic sediment (20%), which indicates that this data may be the result of the addition of various amounts of these two sediment types over time. This amount of sediment mixing is consistent with the results of previous studies (e.g., Tera et al., 1986; Ryan and Langmuir, 1988; Woodhead, 1989; Elliot et al., 1997; Peate and Pearce, 1998; Pearce et al., 2005; Wade et al., 2005; Woodhead et al., 2005;

Chauvel et al., 2009). Ba/La varies with almost constant εNd for the Aleutians and the modern

Marianas and Izu segments, indicating that the Ba (and other fluid mobile elements) concentrations of these arcs are primarily controlled by factors other than the sediment contribution, such as fractionation during dehydration, which will be discussed below.

Based on the r2 values, sediment addition is responsible for 33.3% and 11.9% of the Izu segment’s (La/Sm)N and Ba/La, respectively, but it is not a significant factor for Th/La or K2O.

Sediment addition is responsible for only 21.1% of the variation in (La/Sm)N in the modern

Marianas, 29.7% of Th/La, and 62.4% of K2O, but is not a significant factor for Ba/La. Thus, overall, sediment addition plays a larger role in the incompatible trace element variation in the

Mariana segment than in the Izu segment, which is consistent with the fact that more sediment addition occurred in the Mariana segment (Figure 3.7), judging from the range in εNd. However, as these correlations show, sediment addition appears to account for only a fraction of the incompatible element variations, so other factors must be involved, i.e., differences in the extent of melting and/or dehydration of the slab or extent of melting of the modified mantle wedge, which will be discussed further in the next section.

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Umnak’s Th/La is significantly higher than that of the other Aleutian Islands despite

similar REE enrichments and εNd. This may be due to Umnak’s location farther east and closer to

the continental margin where more terrigenous, Th-rich sediment may be subducting (Kay and

Kay, 1994; Class et al., 2000).

Extent of melting of the mantle wedge at > 30 Ma: To the extent that (La/Sm)N (Figure

3.9) or Th/La (Figure 3.10) does not correlate with εNd, other factors such as the extent of

melting are controlling the incompatible element ratios, such as (La/Sm)N. Larger degrees of

partial melting would produce lower (La/Sm)N at a given εNd.

The extent of melting depends on the mantle potential temperature, the height to which

the mantle can rise once melting begins, and water. The latter is particularly important in

subduction zone volcanism. The water released during subduction is mostly contained within

amphibole, chlorite, and serpentine. Chlorite breaks down at 2-3.6 GPa and 800-820°C, above

the H2O-saturated solidus, so it can result in 2 wt% H2O in bulk peridotite even in hot subduction

zones (Grove et al., 2012). Iwamori (1998) concluded that a thin layer of serpentinite with

chlorite forms as the slab enters the mantle wedge. This layer accommodates all of the H2O

released (8 wt %) and carries the H2O to ~150 km where chlorite and serpentine breakdown. This results in a vertical column of H2O transport. When the H2O reaches 80 km, the temperature

exceeds that of the solidus and melting occurs. Ribeiro and Lee (2017) concluded that

serpentinized forearc mantle that is dragged down by the motion of the descending slab can

provide enough water (7-78% of water injected into trench) to account for the output of H2O

beneath the arc. This explains why the water budgets and geochemical signatures of hot and cold

subduction zones are similar.

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Peacock (1990) determined that the melting of amphibole-bearing oceanic crust only

occurs within a few tens of millions of years of subduction initiation (<50 Ma), i.e., the age of

the Aleutian and IBM arcs. In colder subduction zones (older subduction zones and those built in

older oceanic crust), partial melting occurs primarily in the mantle wedge due to fluid

infiltration, such as from the slab component being carried by low viscosity slab fluids that

ascend to metasomatize and partially melt the mantle wedge.

Grove et al. (2012) states that at >3 GPa, the presence of H2O depresses the peridotite

solidus, causing partial melting. Initial melts are H2O-rich (30 wt%) and are positively buoyant, so they can ascend by porous flow or as diapirs. During this ascent, the hotter melt dissolves the surrounding peridotite until thermal and chemical equilibrium is reached. Grove et al. (2012) conclude that H2O supply is the limiting factor in melting. The H2O content of arc magmas (6-8

wt%) is consistent with the upper limit of pre-eruptive H2O contents (12-14 wt%). The

maximum amount of melt (15 wt%) occurs in the hottest part of the wedge over a range of 40-60

km (Grove et al., 2006). The similarity in mantle water contents is due to the relationship

between the water content and melt fraction. Larger H2O contents lead to larger melt fractions

and larger melt fractions dilute the H2O, so all melts have similar H2O contents. This controls the

viscosity and stalling depth of the melts and leads to a similar range of H2O contents in arc

erupted magmas (e.g., Plank et al., 2013; Grove et al., 2012).

Although partial melting caused by H2O lowering the peridotite solidus is the largest

cause of melting in subduction zones, decompression melting at temperatures above the dry

peridotite solidus also occurs, especially under thin crust (<5%) (Sisson and Bronto, 1998;

Portnyagin et al., 2007).

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Schmidt and Poli (1998) concluded that approximately 15-25% of the initially subducted

H2O is released below volcanic arcs; and the continuous dehydration of hydrous material leads to

H2O being available at 150-200 km. They also point out that there is no difference between the

water contents and fluid release depths of hot and cold subduction zones. Thirty to seventy

percent of the H2O subducted is released to the cold corner of the mantle wedge where serpentine

is stable. They argue that the volcanic front forms above the mantle wedge isotherm (1300°C

isotherm) where the extent of melting is sufficient to allow the extraction of magmas. The

position of his 1300-1350°C isotherm is consistent with the depth of the slab below arcs.

England and Katz (2010) believe that the location of the arc front corresponds to the location

where the boundary defined by the anhydrous solidus is closest to the trench. At this point, the

magma formed at temperatures above the anhydrous solidus gets focused toward the nose of the

region bounded by the anhydrous solidus, and upward flow erodes a path that is continuously

used by ascending magmas, which defines the arc front.

The presence of the extensional upwelling associated with the rifting on northern Attu

combined with subduction under southern Attu resulted in the older Attu magmas being

predominantly produced by decompression melting of mantle that had been influenced by

slab/sediment subduction (see Chapter 1 for a detailed discussion). Similar to those on Attu, the

depleted Kiska samples are the result of an extensional regime on the ridge itself early in the

history of the Aleutian Arc and to a lesser extent in the present day as evidenced by the presence

of summit basins on the ridge.

The forearc of the Mariana segment also includes volcanics with MORB-like Ta/La

(Reagan et al., 2010) similar to the older Attu samples. Reagan et al. (2010) proposed that these

MORB-like tholeiitic forearc basalts are the result of decompression melting that occurred when

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the Pacific Plate first began sinking below the Philippine Plate, which resulted in a near trench

spreading environment with little flux from the slab. Reagan et al. (2010) believe these basalts

may have been very voluminous (similar to the large volume MORB-like magmas on Attu), but

have subsequently been removed by back-arc rifting, although they have been found in back arc

drilling west of the Bonin segment (Arculus et al., 2015) and could be the most abundant in the

fore-arc. These older Attu and Mariana samples, which are depleted in Th/La, La/Sm, and

Sm/Yb, have arrays that begin in the MORB field. The lower end of the (La/Sm)N and (Sm/Yb)N

(Figure 3.5B) ranges of the older Bonin and Mariana arrays overlap the MORB field. These

LREE-depleted IBM samples also have depleted Th/La, Ba/La, and εNd signatures. The Mariana

samples with MORB-like values are the depleted forearc tholeiites, which were erupted before

the depleted mantle could become very enriched by sediment and fluid inputs. These were also

voluminous melts, so the large degrees of melting also contributed to their depletion. Not all of

the older samples plot in the MORB field; other units in the >30 Ma category are less MORB-

like and are discussed in more detail below.

The Bonin segment displays two distinct trends on the (La/Sm)N and (Sm/Yb)N plot

(Figure 3.5B), i.e., the non-boninite Bonin samples (black symbols) and the boninites (yellow symbols), both of which overlap the MORB field at one end. The non-boninite samples display the expected positive correlation between (Sm/Yb)N and (La/Sm)N, while the boninites display

an inverse correlation. The positive La/Sm-Sm/Yb Bonin trend reflects the episode of volcanism

subsequent to the boninite episode demonstrating that the mantle had been recharged with

incompatible elements. The (La/Sm)N, Ba/La, and Th/La of this later event are largely controlled

by the degree of partial melting.

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The negative La/Sm-Sm/Yb trend (Figure 3.5B) is indicative of LREE depletion, which is the result of the melting mechanism that produces boninites. In the IBM Arc, large volume decompression melting produced the forearc basalts for approximately the first 2-4 Ma (Straub et al., 2015; Ishizuka et al., 2018; Brandl et al., 2017; Reagan et al., 2017; Stern and Gerya, 2017).

This caused extensive depletion of the mantle wedge, resulting in a depleted harzburgite source for any further melting. As subduction continued and the slab descended to deeper depths and higher temperatures, slab fluids enriched in fluid mobile elements (e.g., Rb, Ba, K) were released and metasomatized the depleted harzburgite mantle, causing flux melting at shallow depths and lower temperatures, which produces boninites (Stern and Gerya, 2017; Reagan et al., 2009;

Ishizuka et al., 2018; Reagan et al., 2017). This continued for approximately 4 Ma (at ~48-44

Ma) (Straub et al., 2015; Ishizuka et al., 2018; Reagan et al., 2009; Reagan et al., 2017; Brandl et al., 2017). The fluid enrichment that induced melting of this highly depleted mantle resulted in the boninites, which are LREE-depleted, but have a strong fluid mobile trace element signature

(Stern and Gerya, 2017). Reagan et al. (2017) estimates that the addition of 1% water can cause

20% melting of mantle previously depleted by decompression melting. Eventually, new material was advected into the mantle wedge recharging the area, possibly due to foundering of the slab, and the boninite phase of the arc transitioned to more “normal” mature arc volcanism at ~44 Ma,

7-8 Ma after subduction initiation (Ishizuka et al., 2018; Reagan et al., 2009; Brandl et al., 2017).

The range of (La/Sm)N values at a given εNd value indicates that the variation in the degree of partial melting was largest in the Mariana segment and on Attu. In the former, there is also considerable variation in the sediment contribution, while there is almost no sediment contribution to the Attu magmas.

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Extent of melting of the mantle wedge from 0-5 Ma: On the (Sm/Yb)N vs (La/Sm)N and

Th/La vs εNd plots (Figures 3.5A, 3.10), the IBM fields extend to lower values than do the

Aleutians and the IBM has lower on average Th/La, (Sm/Yb)N, and (La/Sm)N ratios compared to

those of the Aleutians, suggesting that on average, the IBM magmas were produced by larger

extents of melting and/or from a source less enriched by ongoing subduction. The Mariana’s

high-K volcanics (not shoshonites; Sun and Stern, 2001), which are also REE enriched, are an

exception and were likely produced by smaller degrees of partial melting. The Mariana segment

displays considerably more variability in (La/Sm)N at a given εNd (Figure 3.9) than the Izu

segment, suggesting a greater variation in the extent of melting in the former. Based on these

same ratios, there appears to have been less variability in the extents of partial melting in the

Aleutians, particularly in the Western Aleutians. Within the Aleutians, extents of melting appear

to have been lowest on Adak, followed by Attu and Kiska.

The effects of crustal thickness: The differences in the crustal thickness of the arcs may affect the fractionation trends and hence the magma compositions. The crustal thickness under

Adak is ~37 km and that under Attu is ~27 km (Janiszewski et al., 2013); while the crustal thickness under the IBM is 20-15 km, north to south (Stern et al., 2003). This difference in thickness is the result of back-arc spreading in the IBM removing the old ridges from the volcanic front instead of a continual build up; whereas, the Aleutian Arc has continued to build on the original ridge (built to ~35-17 km by 35 Ma) (Kay et al., in press), although the volcanic line has shifted to the north (Kay and Kay, 1994; Kay et al., 1990; Jicha and Kay, 2018). The fact that a northward migration of the arc front is observed throughout the Aleutian Arc, while in the

IBM, the arc front has remained stable through time is also related to the combined effects of

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crustal thickening and the tectonic regime (e.g., subduction erosion). As the arc crust thickens

over time, the mantle melt column is truncated and the volume of melt decreases. The location of

melting, and thus the location of arc front magmatism, shifts further from the trench (deeper), as

is seen in the Aleutians. In addition, subduction erosion can erode the upper plate, causing

melting that occurs at the same depth to erupt farther north on the upper plate. Whereas, in the

IBM, crustal thickening is balanced by extension, e.g., back arc spreading, so the overall crustal

thickness is stable, and the arc front does not migrate (Karlstrom et al., 2014; Portnyagin et al.,

2007).

This reduction in crustal thickness could result in lower pressure, shallower melting and a

longer melting column, leading to more melting (Plank and Langmuir, 1988). This would allow

for the generation of boninites when the residual highly depleted mantle melted with a water-rich

fluid flux at shallow levels later in the IBM history (Reagan et al., 2010). These boninites occur

stratigraphically after the voluminous arc melts and concurrent with the initiation of back-arc

spreading and the tholeiitic basalts produced during back-arc basin formation (Crawford et al.,

1981). The lack of this back-arc spreading and crustal thinning in the Aleutians may account for

the lack of boninites in the Aleutians. The more extensive back-arc rifting in the southern part of the IBM results in the thinnest crust, which would also favor larger degrees of melting, as is indicated by the (Sm/Yb)N and (La/Sm)N ratios (Figure 3.5). Crustal thickness also affects the

differentiation paths of magmas. Magmas erupting through thicker crust take longer to reach the

surface and experience more cooling (see review in Farmer and Lee, 2017 and references

therein). Magmas erupting through thick crust exhibit more iron depletion at a given MgO

content, indicating that these melts underwent early saturation of magnetite and differentiated

under more oxidized conditions (see review in Farmer and Lee, 2017 and references therein).

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The thicker crust in the Aleutians could also lead to magmas stagnating in the crust and

fractionating, which could contribute to the higher incompatible element and K2O concentrations

in the Aleutians, and in particular, on Adak (Kay et al., 1990; Kay and Kay, 1994). This would

be especially true in open system refilled-trapped-fractionated (RTF) magma chambers.

Flux of slab-derived aqueous fluids to the arc systems: In addition to variations in sediment contribution and extents of melting, variations in fluid-mobile components such as K

(Figure 3.2) and Ba can reflect the extent to which the mantle wedge is metasomatized by fluids derived from the subducting slab. Portnyagin et al. (2007) concluded that the fluid-rich component leaving the slab is an H2O-rich silicate melt or a supercritical liquid. This fluid rich component transports large amounts of LILE, LREE, and Th. Increasing subduction depths cause the fluid component to transition from a hydrous fluid to a super critical liquid. A temperature of

700-800°C is required to melt water-saturated sediment or basalts, i.e., greater depths. Stolper and Newman (1994) found positive correlations between H2O and Al, Ba, Ca, Cl, K, La, Rb, Sr,

87 86 Sr/ Sr, Th, and U; and negative correlations between H2O and C, Fe, Mn, Nd, P, Pb, Sm, and

Ta. However, they determined that all of these correlations are inherited from the parental

magmas. Higher H2O contents correlate with higher degrees of partial melting, and increasing

the H2O content by 0.2 wt% increases the degree of melting by 12%. Cervantes and Wallace

(2003) determined that H2O enriches magmas in LILEs and LREEs relative to HFSEs. Low

LILEs and LREEs relative to HREEs correlate with low H2O contents, indicating that these

magmas formed largely by decompression melting. They also conclude that the variable Nb and

Ta contents observed are caused by fluid enrichment processes and that the presence of variable

depletions of the mantle wedge is due to earlier partial melting events. Thus, the geochemical

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signatures of the melts are due to variably depleted mantle wedge overprinted by enrichments

from an H2O-rich slab component.

Grove et al. (2002) proposed a fluid rich component (small degree melts of supercritical

hydrous fluids derived from the oceanic crust and sediment) that contributes >99% of the

incompatible contents (Pb, Sr, and εNd), while reactive porous flow and flux melting in the

mantle wedge contribute >90% of the major element contents. Although a significant amount of

Na2O, H2O, and K2O is derived from the slab.

The importance of the sediment contribution can be assessed by the extent to which parameters such as K2O and Ba/La correlate with εNd (Figures 3.8, 3.3), while the importance of

melting can be assessed from the correlation of εNd with fluid-immobile element ratios such as

La/Sm. The fluid-mobile elements are the elements most likely to be disturbed by secondary processes. Rubenstone (1984) and Yogodzinski et al. (1993) concluded that the Ba/La and K2O

(Figures 3.8, 3.2B. 3.3) of the older Aleutian samples have been altered, particularly to higher

values, and cannot be considered to be original values. Alteration also makes the Ba/La and K2O

values questionable in the older IBM data, making it difficult to assess the role of fluid transport

in the ancient arcs. Consequently, the discussion here focuses exclusively on the young/modern

arcs.

The IBM and Aleutian samples on average have similar Ba/La (Figure 3.6A), but the

IBM data ranges to higher and lower values. Based on the r2 values determined in the sediment

subduction section, sediment contribution is a significant control on Ba/La in the Izu segment,

but not in the Aleutians or in the modern Mariana segment. The Izu segment and, to a lesser

extent, the Mariana segment exhibit particularly large Ba/La ranges over limited Th/La ranges

(Figure 3.6A), indicating that their Ba enrichment is independent of the extent of melting. In

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these cases, the Ba enrichment must be due to metasomatism of the mantle wedge by slab- derived fluids. Attu appears to display a similar trend, but there is relatively little data on the modern volcanics on Attu Island due to the fact that there is no active volcanism on the island itself (we do not consider the submarine volcanism north of Attu in this study). In contrast, the

Ba/La of the Central Aleutians (Adak and Umnak) correlates well with Th/La, suggesting that the variation is due to variable extents of melting. Kiska has very few data points for this age, but they do appear to correlate with Th/La, suggesting that its Ba variation is also due to the extent of melting and crystal fractionation.

Sediment input also appears to be an important factor influencing the K2O of the modern

Mariana segment (r2 = 62.4%), but appears to be less important for the Izu segment and the

Aleutians. The modern Mariana and Izu segments and Attu exhibit variable K2O that is

independent of Th/La (Figure 3.3), suggesting that the portion of their K2O enrichment that is not

accounted for by sediment addition is not due to the extent of melting and hence is likely due to

fluid metasomatism. Adak (Adagdak and Moffett) exhibits only a slight increase in K2O with increasing Th/La, suggesting only a small effect on K2O from the extent of melting. Though

Kiska has few data points, it does exhibit a large increase in K2O with increasing Th/La,

suggesting that the extent of melting is important to its K2O enrichment. These trends indicate

that the importance of fluids and metasomatism decreases to the west in the Aleutian Arc.

Chemical trends through time: During the entire history of the Aleutian Arc the influence

of subducted sediment on magma compositions experiences relatively small changes with time

compared to the IBM (Figure 3.12). The degree of partial melting does seem to have decreased

with time (Figure 3.12), especially in the western Aleutians. The western Aleutians (Attu and

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Kiska) experienced a somewhat abrupt change in chemistry, including a shift from large volume

tholeiitic to smaller volume calc-alkaline compositions at ~6 Ma; greater incompatible element enrichment, presumably reflecting smaller extents of melting; and larger Ta and Nb depletions, but it did not experience a shift in the Sr and Nd isotopes around 6 Ma. This change corresponds in time to a change in Pacific Plate motion at 6 Ma (Austermann et al., 2011), which may well be the cause. This change in plate motion had a more significant effect on the more oblique western section of the Aleutian Arc (Attu) than on the central and eastern Aleutians (Adak and Umnak) and the IBM. For further details, see Chapter 1.

The changes in the IBM through time are more significant and include a decrease in the amount of sediment being subducted and a change in sediment composition in the Mariana segment to one richer in volcanoclastics (Peate and Pearce, 1998; Ishikawa and Tera, 1999; Stern et al., 2003; Pearce et al., 2005). Overall, the degree to which the sediment affected the incompatible trace elements decreased. The degree of partial melting in the Mariana segment was largely consistent over the history of the arc, but some small degree melts, the high-K

samples, occurred only in the modern Mariana segment. In summary, as far as one can judge

from the data, which is particularly limited by the lack of data on intermediate age samples (due

to the small number of volcanic events at 10-25 Ma), there is no dramatic shift in chemistry in

the IBM Arc subsequent to the initial episode, but rather a slight change due to changes in

subduction input and extent of melting.

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Figure 3.12: A) age vs. Th/La, B) age vs. εNd, C) age vs. La/Sm [normalized to C1 chondrite (McDonough and Sun, 1995)], and D) age vs. Ta/La. Attu and Kiska data for dated (with error bars) and age estimated samples. Average values of MORB depicted by black lines and error range by blue fields as determined by White and Klein (2014). Data references as in Figure 3.2.

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IMPLICATIONS FOR SUBDUCTION INITIATION AROUND THE PACIFIC RIM

The Aleutian and Izu-Bonin-Mariana arcs have very different beginnings and both

similarities and differences in their evolution and current structures. Though the initiation of both

arcs occurred at ~50 Ma by subduction of the Pacific Plate, their tectonic environments and

initiation mechanisms were quite different. The Aleutian Arc initiated in a compressive

environment to accommodate the motion of the Pacific Plate previously accommodated by

subduction along the Kamchatka-Koryak Margin, which became obstructed by the accretion of

the Olyutorsky Arc shortly before the initiation of Aleutian subduction (see review of Scholl,

2007). In contrast, the IBM initiated with subsidence of the Pacific Plate due to the difference in

density between the older Pacific Plate and the much younger Philippine Sea Plate. This

subsidence (triggered by changes in the motion of the Pacific Plate) eventually transitioned to

true subduction (Arculus et al., 2015).

This initial difference in the subduction zone environments led to a large initial difference in chemistry. The IBM’s initial magmatism included the eruption of depleted tholeiites, boninites, and low-K rhyodacites (Hickey and Frey, 1982; Stern et al., 1991; Taylor et al., 1994;

Stern et al., 2003, Reagan et al., 2010). On Attu, the initial magmatism was arc-like although

with a weak sediment/fluid subduction zone signature, which later transitioned for a time to a

more depleted chemistry due to incipient intra-arc rifting in the north. The compressive environment of the Aleutian subduction initiation resulted in the subduction signature seen in the central and eastern Aleutians today. The extensional subsidence environment of early IBM subduction resulted in the depleted hot mantle being enriched with fluids fluxed from the

subsiding slab, which caused extensive shallow melting of harzburgite, resulting in the eruption

of large volumes of boninites.

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Once true subduction started and the arc massif built up, the chemistry of the IBM gained

a more classic subduction signature. However, the various differences in the geometry of

subduction, the motion of the surrounding plates, the resulting structure of the arc, the age of the

subducting plate, and the subduction of various sediments resulted in differences between the

chemistry of the Aleutian and IBM arcs. At present, for the IBM, the age and convergence angle

of the subducting plate differs along arc as does the type and amount of sediment being

subducted. The most significant difference may be the back-arc spreading (see comparison in

introduction and references therein) in the IBM Arc system. The presence of back-arc spreading in the IBM removes the older arc material, limiting the crustal thickness and providing the opportunity for low pressure melting, resulting in a difference in the flow of the underlying/surrounding mantle. Sun and Stern (2001) indicate that these factors, as well as the opportunity for the melt to equilibrate with the mantle, are key components for shoshonite formation in the Bonin region.

The Aleutian Arc did not, strictly speaking, experience back-arc spreading, though an episode of incipient intra-arc rifting that never progressed to true back-arc spreading and the development of a back arc basin occurred until ~6 Ma on Attu. This extensional regime in the western Aleutians was probably a result of the more oblique subduction in the west that is also evidenced by the extensional nature of the Komandorsky and Bowers Basins. Perhaps in part due to its thicker crust, volcanism in the western Aleutians later transitioned from tholeiitic to calc- alkaline around the time of a change in Pacific Plate motion at 6 Ma. This change in plate motion seems to have had a more significant effect on the more oblique western Aleutian Arc than on the IBM.

251

In the IBM system, the magnitude and composition of the sediment flux changed over time. The modern IBM received a smaller flux of sediment (Figure 3.3), which had a large effect on a mantle previously depleted by the large volume initial arc volcanism, especially in the

Mariana segment. In addition, sediment subducting beneath the IBM became less enriched (Sr and Nd isotopes, La/Sm) with time due to a transition from terrigenous and pelagic sediment to volcanoclastic sediment, particularly in the Mariana segment (Peate and Pearce, 1998; Ishikawa and Tera, 1999; Stern et al., 2003).

In contrast, in the Aleutians there was a transition to a higher, more terrigenous sediment supply with time due to Pleistocene glaciation (Kay and Kay, 1994; Schean et al., 2016; Kay et al., in press), and the more enriched Sr isotopes of the older samples are concluded to be due to seawater alteration (Rubenstone, 1984; Yogodzinski et al., 1993). The sediment flux increases and the composition becomes more enriched along the arc from west to east toward Alaska due to a larger increase in the terrigenous component near the continent.

The tectonic configurations leading to initiation, though different, only seem to control the initial chemistry of the arcs. The subsequent similarities and differences seem to be controlled by the sediment composition and flux, the aqueous fluid flux, the extent of melting, and the influence of the back arc present in the IBM. The initiation of subduction zones around the Pacific at ~50 Ma and the chemistry of the arcs resulting from them clearly show the diversity of environments that can result in subduction and the importance of later subduction factors as opposed to initiation factors in controlling the resulting evolution and chemistry of arcs. The initiation factors seem to be most important to the initial large scale magmatism and are relatively less important to later events.

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261

Yogodozinski, G.M., Brown, S.T., Kelemen, P.B., Vervoort, J.D., Portnyagin, M., Sims, K.W.W., Hoernle, K., Jicha, B.R., Werner, R., 2015, The Role of Subducted Basalt in the Source of Island Arc Magmas: Evidence from Seafloor Lavas of the Western Aleutians: Journal of Petrology, v. 56, p.441-492.

262

APPENDIX A: ANALYTICAL METHODS

The samples were cut, the saw marks were ground off with grinding compound, the slabs

were broken into ~1 cm size pieces, and powdered in an alumina shatterbox in preparation for

major element, trace element, and isotopic analyses.

Major Elements

H2O was determined for samples by weighing the sample before and after heating at

110°C overnight and subsequent cooling in a desiccator; loss on ignition was then subsequently

determined by weighting after heating at 900°C for 30 minutes and subsequent cooling in a

desiccator. Major elements were analyzed by inductively coupled plasma optical emission

spectroscopy (ICP-OES). Undried sample powders (0.1 g) were mixed with 0.4 g of lithium

metaborate and fused at 1050°C for a total of 15 minutes, the samples were swirled after 10

minutes. The molten beads were then added to 50 ml of trace element grade 1.5 N HNO3 and

shaken for 1-2 hours as needed. Immediately after shaking, dilutions of 6 times were performed

to prevent silica gel from forming. This resulted in approximately 3000 times dilution of the

samples. All dilutions were done by mass. Samples were then run with blanks, a check-standard,

and the following standards: AGV-1, BCR-2, BHVO-2, BEN, BIR-1, DNC-1, W-2, GSP-2,

QLO-1, and RGM-1. A check standard (BHVO-2, BCR-2, or W-3) was analyzed periodically throughout the run to measure drift, then used to correct for drift when necessary. Measured average values and standard deviations as well as the acceptable range of values are listed below for each check standard used. Accepted values for BHVO-2 are from Wilson, 1997b; accepted values for BCR-2 are from Wilson, 1997a; and accepted values for W-2 are from Flanagan, and

Gottfried, 1980; Flanagan, 1984; Gladney and Roelandts, 1988; Govindaraju, 1994.

263

BHVO-2 (n = 9) accepted measured value value 1σ SiO 2 49.9±0.6 49.81 0.6798 TiO2 2.73±0.04 2.69 0.5820 Al2O3 13.5±0.2 13.60 0.6996 FeO 12.3±0.2 12.40 1.4570 MnO 0.17±0.005 0.17 0.9118 MgO 7.23±0.12 7.45 0.7470 CaO 11.4±0.2 11.58 1.2470 Na2O 2.22±0.08 2.19 0.8198 K2O 0.52±0.01 0.49 1.0192 P2O5 0.27±0.02 0.27 1.1405

BCR-2 (n = 28) accepted measured value value 1σ SiO 2 54.1±0.8 54.12 1.7480 TiO2 2.26±0.05 2.23 1.8134 Al2O3 13.5±0.2 13.58 1.6302 FeO 13.8±0.2 13.79 1.5887 MnO 0.196±0.008 0.20 1.4238 MgO 3.59±0.05 3.67 1.6269 CaO 7.12±0.11 7.17 1.3455 Na2O 3.16±0.11 3.13 1.7285 K2O 1.79±0.05 1.78 1.3176 P2O5 0.35±0.02 0.36 1.7682

W-2 (n = 15) accepted measured value value % error SiO 2 52.68±0.29 52.04 0.9151 TiO2 1.06±0.01 1.03 1.2680 Al2O3 15.45±0.16 15.37 1.2308 FeO 10.83±0.21 10.66 0.8835 MnO 0.167±0.004 0.17 1.0023 MgO 6.37±0.058 6.58 1.2807 CaO 10.86±0.078 10.90 0.6909 Na2O 2.2±0.037 2.19 1.5832 K2O 0.626±0.012 0.61 1.4638 P2O5 0.14±0.12 0.13 0.9931

264

Isotope Ratios Sr and Nd isotope ratios were determined on a VG Sector 54 thermal ionization mass

spectrometer (TIMS) in dynamic multicollector mode. 0.05 g of sample powder was leached in 6

N HCl for 1 hour in Savillex™ capsules, rinsed in milli-Q water, and dried. Four ml of HF and

~0.1 ml perchloric acid were then added, the capsule was capped, and the capsule was heated on low temperature overnight. The solution was then evaporated to dryness, the sample was redissolved twice in 3 ml of HCl, and then the sample was dried. The sample was then dissolved in 1.25 ml of 2.5 N HCl and transferred to an acid washed centrifuge capsule. Samples were centrifuged for 15 minutes at 5000 rpm immediately before being added to ion exchange columns consisting of 50W-AGx8 200-400 mesh resin. After 19 or 25 ml (depending on column

size) of 2.5 N HCl was added to elute the major elements, Sr was eluted with 9 or 14 ml of 2.5 N

HCl. Following this, 3 or 4 ml of 6 N HCl was added to the column, and following that the rare

earth elements were eluted with 12 or 17 ml of 6 N HCl. The Sr and rare earth fractions were

dried overnight. The rare earth fraction was then dissolved in 0.2 ml of 0.23 N HCl. The solution

was added to premade 2 ml columns of Eichrom 100-150 micron Ln™ resin. After washing with

9 ml of 0.23 N HCl, the Nd was eluted with 10 ml of 0.23 N HCl.

Sr was loaded in 2.5 N HCl with the addition of 1μl of TaF onto degassed tungsten

filaments. The samples and standard NBS 987 were analyzed at 1.5-3 x 10-11 A of ion intensity

of 88Sr; typically 90-285 individual ratios were measured and averaged for a single analysis. Nd

was loaded in a 2.5 N HCl and 0.5 M H3PO4 mixture onto degassed Re filaments following

drying of ~0.5 μl of AG50W-x12 resin beads (which served as a reductant). Samples and the

AMES neodymium standard were run at 0.3 x 10-11 A ion intensity of 144Nd; typically 90-300

individual ratios were measured and averaged for a single analysis.

265

NBS 987 was used as the standard for Sr analysis. For n=48, the average value of

87Sr/86Sr was 0.710227 with a standard deviation of 0.000017 and a 2SEM of 0.000008. The accepted 87Sr/86Sr value for NBS 987 is 0.71034 ± 0.00026, which makes our values within the acceptable range. To correct for fractionation, we used a 86Sr/88Sr normalization values of

0.11940. AMES was used as the standard for the Nd analyses. For n=42, the average value of

143Nd/144Nd was 0.512132 with a standard deviation of 0.000008 and a 2σ of 0.000012. The accepted 143Nd/144Nd value for AMES is 0.512131, which makes our values within an acceptable range. Averages of standard analyses for each run were used as the basis for any corrections needed to the measured sample values. To correct for fractionation, we used a 146Nd/144Nd normalization value of 0.1936.

Trace Elements

Trace elements were analyzed by inductively coupled plasma mass spectrometry (ICP-

MS) at Cornell University using a FISONS Element™ instrument and at Colgate University using a Varian instrument. Samples analyzed at Cornell were prepared by dissolving 0.1 g of sample powder in 12 ml of trace element grade HNO3 and 6 ml of HF in Savillex™ capsules and heated overnight. The solution was dried, redissolved in 4.5 ml HNO3, and heated in the same capsules overnight. Seventeen ml of milli-Q water was then added, the solution was shaken, heated for 10 minutes, and allowed to cool. The samples were then centrifuged for 10 minutes at

3000 rpm. After centrifuging, the solution was diluted by 100 times and then by another 10 times to make a 1000 times dilution. This resulted in overall dilutions of approximately 20,000 and 200,000 times respectively. All of the dilutions were done by mass. The samples were run in triplicate with blanks, a check standard, and the following standards: PAL, AGV-1, BCR-2,

266

BHVO-2, BEN, BIR-1, DNC-1, W-2, G-2, GSP-1, and QLO-1. The check standard used was

BEN. See the table below for the average measured value and the standard deviation as well as the accepted values.

Samples analyzed at Colgate were prepared by dissolving 0.1 g of sample powder in 7.2 ml of trace element grade HNO3 and 4.8 ml of HF in Savillex™ capsules and heated for 24

hours. High silica samples were heated in Teflon capsules inside a metal “bomb” at 225°C for 24

hrs, were allowed to cool to room temperature, and were removed from the bombs before the

drying step. This solution was then dried, redissolved in 12 ml of 50:50 HNO3 and milli-Q water,

and heated in capsules for approximately 48 hours. After the beakers cooled, the solution was

transferred to a 125 ml acid washed bottle and diluted with 88 ml of milli-Q water. These

samples are then run with a check standard and the following standards: BIR-1, W-2, DNC-1,

BHVO-2, AGV-2, and AGV-2 (diluted by half). The unknown used was W-2. See the table

below for the average measured value and the standard deviation as well as the accepted values.

Accepted values for W-2 are from Flanagan, and Gottfried, 1980; Flanagan, 1984; Gladney and

Roelandts, 1988; Govindaraju, 1994; Accepted values for Ben are from Govindaraju, 1989.

267

BEN (n = 22) measured accepted value 1σ value Sc 28.76 11.58 22.0±1.5 V 232.3 2.221 235±10 Cr 310.7 1.883 360±12 Co 57.28 1.269 60.0±2.0 Ni 241.6 1.264 267±7.0 Cu 79.60 5.448 72±3 Zn 129.0 3.697 120±13 Rb 53.96 3.306 47±2 Sr 1387 5.542 1370±25 Y 34.16 0.827 30±1.5 Zr 248.3 4.461 260±10 Nb 100.9 0.909 105±8 Cs 0.858 0.881 0.8±0.10 Ba 965.7 5.660 1025±30 La 91.55 2.438 82±2 Ce 154.8 3.045 152±4 Pr 19.31 2.869 18±1 Nd 67.51 2.654 67±2 Sm 14.13 3.604 12.2±0.3 Eu 3.916 2.162 3.60±0.18 Gd 12.15 2.080 10±0.6 Tb 1.613 2.172 1.30±0.10 Dy 7.461 3.307 6.4±0.20 Ho 1.278 1.510 1.1±0.13 Er 3.012 1.655 2.5±0.10 Tm 0.408 1.430 0.34±0.40 Yb 2.627 1.544 1.8±0.2 Lu 0.462 3.645 0.240±0.030 Hf 6.086 3.491 5.6±0.2 Ta 5.739 2.891 5.70±0.400 Pb 4.722 2.848 4±2 Th 15.80 2.070 10±1 U 2.714 2.485 2.40±0.18 Analyzed at Cornell University

268

W-2 (n = 29) measured accepted value 1σ value Sc 40.73 8.425 36±1.1 V 286.1 9.272 260±12 Cr 100.9 6.722 92±4.4 Co 48.16 7.202 43±2.1 Ni 78.13 7.074 70±2.5 Cu 111.6 5.316 110±4.9 Zn 78.53 7.724 80±2 Rb 19.77 2.939 21±1.1 Sr 202.6 4.008 190±3 Y 23.47 2.989 23±1.6 Zr 92.88 5.141 92±4 Nb 8.123 3.525 7.76 Ba 157.6 3.573 171.6±11 La 10.30 2.746 10.07±0.59 Ce 21.57 5.145 22.79±1.5 Pr 2.992 2.574 3.04 Nd 12.57 2.788 12.9±1 Sm 3.258 3.071 3.24±0.13 Eu 1.108 3.776 1.1±0.06 Gd 3.717 2.103 3.73 Tb 0.614 2.233 0.63 Dy 3.712 2.733 3.83±0.8 Ho 0.784 3.431 0.8 Er 2.188 4.056 2.17 Tm 0.316 3.859 0.33 Yb 1.945 3.550 1.98±0.2 Lu 0.289 3.264 0.3 Hf 2.312 2.578 2.3±0.18 Ta 0.498 2.495 0.5 Pb 7.756 1.203 7.81 Th 2.175 6.134 2.21±0.1 U 0.499 4.041 0.497 Analyzed at Colgate University

269

Ar/Ar Geochronology

All Geochronology was done at the WiscAr Geochronology Laboratory in the

Department of Geosciences at the University of Wisconsin-Madison on a MAP 215-50 mass spectrometer. Attu samples included groundmass, amphibole, plagioclase, and albite separates;

Kiska samples included plagioclase and groundmass separates; and Rat Island samples included groundmass and plagioclase separates. Preparation of samples included: crushing and sieving to

250-500 μm, magnetic and dense liquid separation (methylene iodide), and cleaning ultrasonically and by leaching. Unaltered pieces were hand-picked and irradiated for 7, 40, or 50 hours at the Oregon State University TRIGA reactor. The 28.201 Ma Fish Canyon Sanidine was used as the standard. Five 5 mg of sample were heated with a laser for the Eocene samples and

150 mg of sample were incrementally heated with a furnace for the Miocene samples. For further details see (Schean et al., 2016; Jicha et al., 2012)

270

APPENDIX B: ATTU ISLAND SAMPLE DESCRIPTIONS

Table B1: Attu Island Sample Descriptions Sample Series placement of Latitude Longitude Location Yogodzinski, 1993 determination AT57 albite granite suite 53°27.04N 172°58.81E map location estimation AT17 albite granite suite 52°58.62'N 172°55.40'E AT16 AT140B albite granite suite 52°56.15'N 173°01.30'E AT143 AT53 albite granite suite 53°27.53N 173°3.79E map location estimation AT12 albite granite suite 52°58.60'N 172°56.00'E AT11 AT78 albite granite suite 53°17.75N 172°58.10E map location estimation AT24 albite granite suite 52°58.80'N 172°57.60'E AT27, 29 AT25 albite granite suite 52°58.80'N 172°57.60'E AT27, 29 AT77 albite granite suite 53°17.75N 172°58.10E map location estimation AT33 albite granite suite 53°27.53N 172°58.63E map location estimation AT45 albite granite suite 53°29.15N 172°53.74E map location estimation AT8 late dikes 52°58.52'N 172°56.35'E AT6 AT135 late dikes 52°58.80'N 172°57.60'E AT27, 29 AT134 late dikes 52°58.80'N 172°57.60'E AT27, 29 Note: Latitudes and longitudes were estimated from locations based on field maps, listed in the table as “map location estimation”, or were assigned the location of nearby samples, the numbers of which are listed in the table.

271

m (ft)

Elevation 0 0 0 0 0 671( 2200) 701 (2300) 0 640 (2100) 335 (1100) 533 (1750) 457 (1500) 0 0 0 0

Longitude 173°07.88'E 173°16.10'E 173°07.88'E 173°16.21'E 173°08.40'E 173°00.10'E 173°01.75'E 173°10.05'E 172°55.30'E 172°59.75'E 173°00.45'E 173°00.35'E 173°10.83'E 173°10.44'E 173°08.67'E 173°10.91'E

Latitude 52°56.14'N 52°52.96'N 52°56.14'N 52°52.91'N 52°56.22'N 52°57.95'N 52°58.32'N 52°55.04'N 52°51.20'N 52°55.45'N 52°50.10'N 52°50.10'N 52°49.08'N 52°48.40'N 52°47.82'N 52°49.59'N

olv -

cpx -

cpx ol ol cpx cpx ol ------

pl pl pl aphyric aphyric aphyric pl aphyric cpx megacrysts with pl aphyric pl Phenocrysts pl ol aphyric ol pl

basalt basalt basalt basalt basalt basalt basalt basalt andesite basaltic basalt basalt andesite basaltic basalt basalt basalt basalt Rock Type

Structure lava pillow lava pillow lava pillow lava pillow lava pillow lava pillow lava pillow lava pillow diabase massive diabase massive diabase massive diabase massive lava pillow lava pillow lava pillow lava pillow

CP916C MP2 MP10 AT114 AT144 AT101 AT100 CP98 AT73 H962A Sample HO925B HO914B H922 AT61 SB1A SB5A Table B2: Attu Island Sample Descriptions Sample Island Attu B2: Table

272

Elevation m (ft) 0 0 0 0 0 472 ( 1550) 0 0 76 (250) 0 457 (1500) 472 (1550) 76 (250) 640 (2100) vicinity of Alexai Pass 0

172°56.00'E 172°54.60'E 172°55.40'E 172°57.60'E 172°57.45'E 172°54.70'E 172°56.35'E 172°57.45'E 173°01.10'E 172°57.60'E 172°55.00'E 172°56.25'E 172°53.00'E 173°01.30'E 173° 18.5'E 172°17.50'E Longitude

52°59.65'N 52°58.62'N 52°58.80'N 52°58.55'N 52°57.60'N 52°58.52'N 52°58.55'N 52°59.40'N 52°58.80'N 52°57.85'N 52°57.60'N 52°56.65'N 52°56.15'N 52° 51.04'N Latitude 52°58.60'N 52°49.80'N

qtz ox.

- ox - amph - biot - rich rich ox - - -

ox olv olv ox - cpx - - -

amph - -

- qtz. amph amph qtz ox qtz ox amph ox cpx cpx cpx cpx cpx ------amph pl pl pl pl pl pl pl pl pl aphyric with aphyric with amphtrace pl aphyric pl pl pl Phenocrysts pl

granite

albite granite albite rhyolite andesite andesite Rock Type basalt basalt basalt basalt andesite basaltic andesite gabbro rock plutonic rhyolite rhyolite rhyolite albite

from Yogodzinski et al., 1993. al., et from Yogodzinski

Structure dike dike dike dike dike dike stock stock dike lava lava stock stock sill intruding Formation Faneto stock dike

Table B2 (Continued): B2 Table Sample AT11 AT50 AT16 AT27 AT22 AT21 AT6 AT23 AT65 AT29 AT19 AT3 AT56 AT143 AT8032 AT83 Note: All data in this table is is table this in data All Note:

273

173°16.02E 173°16.29E 173°16.06E 173°17.84E 173°17.84E 173°15.09E 173°14.95E 173°13.35E 173°13.04E 173°16.02E 173°17.22E 173°9.66E Longitude

Latitude 53°7.97N 53°7.97N 53°7.97N 53°10.58N 53°11.15N 53°18.24N 53°19.38N 53°17.75N 53°17.42N 53°17.26N 53°15.80N 53°20.35N

mass; mass;

fine intersertal ab, pink pink ab, intersertal fine

aphyric; fine diabase gmass; amyg ch, ch, amyg gmass; diabase fine aphyric; cc sp porph pl, skeletal opq; subophitic subophitic opq; skeletal pl, sp porph ch, ep, masses of qz anhedral gmass; Petrographic Description Petrographic replaced partially pl phen sp porph, gmass hyalopilitic by ch+qz(?), radiating (some pl phen sp porph, gmass hyalopilitic aggragetes); v sp porph, cpx, opq by ch+qz, replaced pl phen sp porph, spinel; red by ch(?), ol replaced gmass hyalopilitic v amyg gmass; hyalopilitic aphyric; qz, ep cc, v amyg gmass; hyalopilitic aphyric; ep, pu qz, cc, g subophitic pl; phen sp porph, ch of masses anhedral gmass; intersertal cpx; pl, phen glom, cc sp amyg cpx, subophytic pl, holocrystalline, ol opq ch), by (replaced hyalopilitic pl; cloudy phen sp porph, gmass; sp ch, amyg veinpu (?), of cc

- 1A

15 - 70B

15 -

16 -

1.6 km N of SB80 1.6 kmof N head form 4 km CH, shore NW km N 4 Ridge, end ofN Fishhook ofof CH head NE point between CH and CH between point NE HB - of E SB80 2 km shore N SB, - SB80 of E 1.6 km shore N SB, Location Description Location SB of head of corner NW of just S Ridge, end ofN Fishhook CH9 of CH9 1.3 km E CH9 of E km 3.2 Black's Beach, 70B - SB80 9 kmof N N shore HB, 6.9 km HO9 NW of - HO9 overlying immediately 25B, 2A

Pillow Lava Pillow Lava Dike Dike Structure Lava Pillow Pillow Lava Flow Pillow Breccia large of Interior lava pillow Pillow Lava Sill Pillow Breccia

1A 5A 6A 18 20 2D - - - - - 8 11 16 18 43D 52A 57A ------Table B3: Attu Island Sample Descriptions Sample Island Attu B3: Table Sample SB80 SB80 SB80 CH9 SB80 SB80 CH9 CH9 CH9 CH9 CH9 CH9 HO9

274

Longitude 173°10.28E 173°9.30E 173°8.42E 173°8.15E 173°8.06E 173°9.66E 173°9.30E 173°10.19E 173°10.68E 173°10.82E 173°10.82E

Latitude 53°22.89N 53°19.38N 53°18.00N 53°18.00N 53°17.75N 53°15.80N 53°16.44N 53°14.16N 53°55.92N 53°56.24N 53°56.24N

s of cc

Petrographic Description Petrographic cpx; clear pl, cloudy phen glom, gmass hyalopilitic hyalopilitic ch; by replaced ol sp porph gmass gmass fine pilotaxitic aphyric; replaced ol cpx, pl, cloudy phen sp porph, ch cc, amyg gmass; hyalopilitc by cc; gmass pilotaxitic pl; clear porph, phen hyalopilitic pl; cloudy phen sp porph, ch ves sp gmass; cc; by replaced ol cpx, pl, porph, phen ch cc, amyg gmass; pilotaxitic ch, pr opq; cpx, pl, holocrystalline, cc; by replaced ol pl, v sp porph vein gmass; hyalopilitic gmass; hyalopilitic cpx; pl, porph, phen cc ch, sp amyg cpx; pl, glom pl, phen sp porph, (?) ch, pr qz, amyg gmass; hyalopilitic

43,

km NW od km NW 61A

25B 25B 25B 25B 25B

- - - - - side of Casco Point, 1 km N of N 1 km Point, of side Casco E side of Casco Point, 3.2 km N N 3.2 km Point, of sideE Casco of end S E S end Location Description Location of NW 10 km of shore HB, N HO9 SE corner of South Arm ofHB, - HO9 overlying N shore of HB, 10.2 km NW of of NW km 10.2 of shore HB, N HO9 3.9 of shore HB, N HO9 of m NW 400 of shore HB, N HO9 of m NW 300 of shore HB, N HO9 HB of Arm North of corner NW - HO9 underlying immediately NE end of middle ofpoint HB N shore of middle point HB of

Pillow Lava Pillow Lava Pillow Breccia Structure Pillow Lava Pillow Lava Pillow Lava Pillow Lava Pillow Breccia Stock Brecciated Lava Pillow Breccia

62A 8 9 14B 22 23C 25B 42 49B 8 16B 16C ------Table B3 (Continued): Sample HO9 HO9 HO9 HO9 HO9 HO9 CP9 CP9 HO9 HO9 HO9 CP9

275

one, one, Longitude 173°10.28E 173°10.42E 173°10.28E 173°5.64 173°9.13E

Latitude 53°53.80N 53°52.82N 53°52.33N 52°48.71 53°55.75N

locations based on field maps. The location of of location The maps. field on based locations

Petrographic Description Petrographic by replaced cpx, ol pl, phen v porph, brecciated v gmass; intersertal ch; hyalopilitic cpx; pl, phen v porph, cc veins gmass; by ch, cc replaced ol pl, porph, phen holocrystalline, intersertal, pl, intersertal, holocrystalline, cpx,opq; ch, cc intersertal pl; cloudy phen sp porph, gmass; anhedral qz, ab, ep, pr

15, were estimated from from estimated 15, were

N shore Massacre Bay, 9.6 km from km from 9.6 Bay, Massacre shoreN Head Point of shore Alexai S Location Description Location Point Murder W of 4.8 km shore, S Point Murder of W 6 km shore, S 1 km SW of S end of Casco Cove Casco end of of S 1 km SW

Structure Pillow Breccia Pillow Lava Pillow Lava Sill? Brecciated Brecciated Lava

12 Rubenst from are longitudes, and latitudes except data, All location. map field the using Earth from Google determined was 15 15 24 2 8 10 ------

MP80 Table B3 (Continued): B3 Table Sample MP80 MP80 MB80 MB80 except samples, all for longitude and latitude The Note: MB80 MB80 MB80 1984.

276

Longitude 173°8.06E 173°7.35E 173°19.16E 173°18.91E 173°19.44E 173°20.51E 173°18.01E 173°18.09E 173°6.55E

Latitude 52°53.31N 52°48.06N 52°49.70N 52°58.20N 52°49.72N 52°58.68N 52°50.66N 52°51.17N 52°52.98N

Location Bedard Cove Bedard Cove Bedard Cove Bedard Cove Sonoma Beach

tz, calcite, calcite, tz,

Petrographic Description Petrographic phen, cpx, opaques; plag, calcite, chlorite, secondary, mineral, needlelike small cpx replacing biotite, opaques cpx, opaques; phen, plag, qtz, calcite, chlorite, secondary, mineral, needlelike small cpx replacing opaques cpx, opaques; phen, plag, q chlorite, secondary, mineral, needlelike small cpx replacing opaques secondary, opaques; phen, plag, small qtz, calcite, chlorite, mineral needlelike opaques; cpx, Phen plag, biotite chlorite, qtz, secondary,

eration Description holocrystalline, inequigranular, alteration slight holocrystalline, inequigranular, subophitic shows fairly is texture, altered holocrystalline, inequigranular, altered moderatly holocrystalline, inequigranular, altered quite holocrystalline, subophitic shows texture, moderate alt

Rock Type Gabbro Diabase Gabbro Diabase Gabbro

Structure fine grained medium grained fine grained, dike? medium grained fine to medium grained

20* 22 27 27A* 28 30B* 32 36 43* ------Table B4: Attu Island Sample Descriptions Sample Island Attu B4: Table Sample AT80 AT80 AT80 AT80 AT80 AT80 AT80 AT80 AT80

277

Longitude 173°6.55E 173°6.07E 173°6.28E 173°15.42E 173°10.39E 173°6.59E 173°6.25E 173°5.63E 173°23.71E

tions

Latitude 52°52.98N 52°47.93N 52°54.28N 52°51.35N 52°48.30N 52°48.63N 52°48.91N 52°48.70N 52°55.26N

Location Sonoma Beach Sonoma Beach Sonoma Beach Jackass Pass Murder Point South of South Pass South of South Pass South of South Pass

ques; chlorite chlorite is ques;

Petrographic Description Petrographic opaques; cpx, Phen plag, biotite, calcite, chlorite, secondary, are that grains needlelike small and ubiquitous cpx, opaques; phen, plag, biotite, qtz, chlorite, secondary, cpx the replacing with opaques cpx, opaques; phen, plag, and calcite, qtz, chlorite, secondary, opaques some biotite, with cpx replacing cpx, opa phen, plag, mineral alteration dominant the cpx, opaques; phen, plag, and chlorite calcite, secondary, biotite cpx, opaques; phen, plag, secondary chlorite biotite and qtz cpx, opaques; phen, plag, with calcite, chlorite, secondary, cpx replacing opaques cpx, opaques; phen, plag, qtz biotite, chlorite, secondary,

Description holocrystalline, texture, subophitic altered highly holocrystalline, texture subophitic holocrystalline, extreme shows alteration holocrystalline, altered highly holocrystalline, moderately altered holocrystalline, amount fair shows holocrystalline, altered highly holocrystalline, moderately altered

Rock Type layered gabbro plutonic Gabbro Gabbro Gabbro Gabbro Gabbro

Structure fine grained block float medium grained fine grained medium grained fine grained medium grained coarse grained coarse grained

48 54* 71 76 78 89* ------87 -

Table B4 (Continued): B4 Table Sample AT80 43A* AT80 AT80 AT80 61A AT80 AT80 AT80 AT8 AT80 Note: data, All estimated isexcept latitude and longitude, form were for1986. * indicates Shelton, samples the latitude which and longitude mapson field based from locations loca map field using Earth from Google determined were other samples all of locations The

278

Longitude 173°8.06E 173°7.35E 173°19.16E 173°18.91E 173°19.44E 173°20.51E 173°18.01E 173°18.09E 173°6.55E 173°6.55E

Latitude 52°53.31N 52°48.06N 52°49.70N 52°58.20N 52°49.72N 52°58.68N 52°50.66N 52°51.17N 52°52.98N 52°52.98N

Location Bedard Cove Bedard Cove Bedard Cove Bedard Cove Sonoma Beach Sonoma Beach

small

opaques; opaques; Petrographic Description Petrographic cpx, opaques; phen, plag, calcite, chlorite, secondary, mineral, needlelike small cpx replacing biotite, opaques phen, cpx, plag, qtz, calcite, chlorite, secondary, mineral, needlelike small cpx replacing opaques cpx, opaques; phen, plag, calcite, qtz, chlorite, secondary, mineral, needlelike small cpx replacing opaques secondary, opaques; phen, plag, qtz, calcite, chlorite, mineral needlelike opaques; cpx, Phen plag, biotite chlorite, qtz, secondary, opaques; cpx, Phen plag, calcite, chlorite, secondary, needlelike small and biotite, ubiquitous that are grains

alteration

Description holocrystalline, inequigranular, slight holocrystalline, inequigranular, subophitic shows fairly is texture, altered holocrystalline, inequigranular, altered moderatly holocrystalline, inequigranular, altered quite holocrystalline, subophitic shows texture, moderate alteratio holocrystalline, subophitic highly texture, altered

Rock Type Gabbro Diabase Gabbro Diabase Gabbro

Structure fine grained medium grained fine grained, dike? medium grained fine to medium grained fine grained block float

20* 22 27 27A* 28 30B* 32 36 43* 43A* ------Island Sample TableDescriptions B5: Sample Attu Island Sample AT80 AT80 AT80 AT80 AT80 AT80 AT80 AT80 AT80 AT80

279

APPENDIX C: RAT ISLAND AND KISKA ISLAND THIN SECTION DESCRIPTIONS

Rat Island RAT-13-1: Aphyric with matrix composed of 95% glass and 5% olivine altering to iddingsite, oxides, and minor plagioclase.

RAT-13-2: 90% matrix and 10% phenocrysts. Matrix is composed of 95% glass and 5% plagioclase and olivine. Phenocrysts are olivine and slightly degraded plagioclase laths.

RAT-13-2B: 30% Matrix and 70% phenocrysts. Matrix is composed of 95% glass and 5% oxides and a few plagioclase. Phenocrysts are 50/50 slightly degraded amphibole and plagioclase with slight zoning (equant and glomerocrysts).

RAT-13-2C: 50% matrix and 50 % phenocrysts. Matrix is composed of 95% glass and 5% plagioclase and minor oxides. Phenocrysts are 60/40 plagioclase and amphibole, with minor orthopyroxene. The plagioclase displays some zoning and all phenocrysts have fair degradation.

RAT-13-3: 80% matrix and 20% phenocrysts. Matrix is composed of 95% glass and 5% iddingsite and minor oxides and plagioclase. Phenocrysts are plagioclase (small equant and laths), very minor clinopyroxene and olivine altered to iddingsite. All phenocrysts have minor degradation.

RAT-13-4: 90% matrix and 10% phenocrysts. Matrix is composed of plagioclase, oxides, glass, and iddingsite. Phenocrysts are slightly degraded plagioclase (laths, equant, and glomerocryst forms), minor clinopyroxene with twinning, and minor olivine altering to iddingsite.

RAT-13-5: 25% matrix and 75% phenocrysts. Matrix is composed of 75% glass and 25% plagioclase and minor oxides. Phenocrysts are 75/25 Plagioclase and amphibole. The amphibole is slightly degraded and the plagioclase is present in equant grains and glomerocrysts with some zoning and degradation.

RAT-13-6: 20% matrix and 80% phenocrysts. Matrix is composed of glass, plagioclase, and oxides. Phenocrysts are 60/40 amphibole and plagioclase. The amphibole is almost isotropic, likely due to oxide replacement, while the plagioclase is present in heavily degraded equant grains and glomerocrysts.

RAT-13-7: 15% matrix and 85% phenocrysts. Matrix is composed of 90% glass and 10% plagioclase. Phenocrysts are grey-black amphibole with oxide replacement; heavily degraded plagioclase present as large crystals and small equant grains, with no twinning evident; minor oxides, and greyish amorphous shapes and veins, which were likely deposited in vesicles and cracks during weathering.

RAT-13-8: 40% matrix and 60% phenocrysts. Matrix is composed of glass, fairly large amphibole and plagioclase, and minor small oxides. Phenocrysts are 50/50 amphibole and equant plagioclase and a few larger oxides.

280

Kiska Island KS-12-1: 50% matrix and 50% phenocrysts. Matrix is composed of 98% glass and a few olivine and plagioclase. Phenocrysts are plagioclase laths, some with zoning or core degradation, and minor olivine, oxides and clinopyroxene.

KS-12-3: 20% matrix and 80% phenocrysts. Matrix composed of 50% glass, 40% plagioclase, and 10% olivine. Phenocrysts are plagioclase, equant grains and glomerocrysts, with some zoning and heavy alteration; chlorite filling vesicles; minor olivine, less altered than the plagioclase; and very minor oxides.

KS-12-4: 15% matrix and 85% phenocrysts. Matrix is composed of 40% glass, 40% oxides, and 20% plagioclase. Phenocrysts are heavily fractured and degraded plagioclase (equant and laths), minor olivine, oxides, and clinopyroxene, with twinning. Also present is secondary quartz filling fractures and vesicles, and some very heavily altered glomerocrysts that were probably plagioclase originally.

KS-12-5: 50% matrix and 50% phenocrysts. Matrix is composed of chlorite, oxides, plagioclase laths, and minor amounts of very altered olivine. Phenocrysts are heavily altered plagioclase equant grains and glomerocrysts and chlorite filling veins and vesicles.

KS-12-7: Aphyric, with matrix composed of glass, plagioclase laths, oxides, and minor olivine.

KS-12-7N: Aphyric, with matrix composed of glass, plagioclase laths, and oxides. Alteration includes vesicles and veins filled with secondary quartz and a brown fibrous mineral. One very small olivine was also observed.

KS-12-8: 30% matrix and 70% phenocrysts. Matrix is composed of glass, plagioclase laths, olivine, and oxides. Phenocrysts are 50% zoned plagioclase (equant, elongate, and forms), some with core degradation, and minor olivine.

KS-12-9: 80% matrix and 20% phenocrysts. Matrix is composed of glass, plagioclase laths, and oxides. Phenocrysts are equant grains and laths of plagioclase and minor needles and crystals of amphibole.

KS-12-10: 15% matrix and 85% phenocrysts. Matrix is composed of olivine, equant plagioclase, and oxides. Phenocrysts are fairly unaltered equant grains and laths of plagioclase, more altered and fractured and less abundant olivine, and minor oxides, which occur primarily in or near olivine.

KS-12-12: 15% matrix and 85% phenocrysts. Matrix is composed of glass, olivine, oxides, and equant plagioclase. Phenocrysts are olivine completely altered to iddingsite with distinct olivine shape, a few fresher olivine, and equant grains and laths of plagioclase with a few fairly unaltered and zoned and others core degradation,

KS-12-13: Aphyric, with matrix composed of glass, plagioclase laths, iddingsite, and oxides. The vesicles have been filled by a fibrous secondary mineral.

281

KS-12-14: 75% matrix and 25% phenocrysts. Matrix is composed of glass, plagioclase laths, and minor olivine and oxides. Phenocrysts are plagioclase, with slight degradation and some zoning, in the form of equant grains and fairly unaltered smaller laths; minor olivine and clinopyroxene, and a few oxides. Vesicles are present but have not been filled by secondary mineralization.

KS-12-15: 30% matrix and 70% phenocrysts. Matrix is composed of chlorite and oxides. Phenocrysts are heavily altered plagioclase (equant, laths, and glomerocrysts forms) with chlorite ingrowth, completely altered olivine that retains its shape, and chlorite.

KS-12-16: 20% matrix and 80% phenocrysts. Matrix is composed of olivine, equant plagioclase, oxides, and some glass. Phenocrysts are equant grains and laths of plagioclase, with minor zoning and slight degradation, and more altered olivine.

KS-12-17: Phenocrysts are cpx with some twinning, olivine, more altered plagioclase laths and glomerocrysts (glomerocrysts are the most altered), and iddingsite.

KS-12-18: Phenocrysts are groups of smaller, heavily altered plagioclase laths; clinopyroxene, with twinning common; minor olivine, and iddingsite and chlorite from alteration.

KS-12-20: 85% matrix and 15% phenocrysts. Matrix is composed of chlorite, other alteration products, and possibly some quartz. Phenocrysts are extremely altered plagioclase (equant) and quartz.

282

APPENDIX D: ADDITIONAL DATA

Table D1: Attu Island Major Element Data AT3‡ AT5‡ AT6‡ AT11‡ AT12‡ AT16‡ AT17‡ AT19‡ AT21‡

SiO2 74.51 70.88 47.48 49.9 69.53 48.7 67.91 73.79 58.35

TiO2 0.46 0.58 4.44 2.16 0.77 2.95 0.86 0.29 2.04

Al2O3 12.93 14.82 15.35 15.5 14.89 15.58 15.45 13.79 16.24 FeO 2.12 3.57 13.3 10.42 3.37 11.78 4.22 2.71 8.51 MnO 0.03 0.06 0.25 0.17 0.05 0.21 0.06 0.04 0.12 MgO 0.3 0.71 4.91 6.57 0.63 7.5 1.07 0.24 3.67 CaO 0.38 1.29 10.21 8.6 1.49 7.33 1.67 0.39 3.94

Na2O 6.43 6.09 3.35 3.84 8.13 4.32 8.04 6.84 6.12

K2O 2.64 2.02 0.54 1.57 0.16 0.53 1.01 1.61 1.87

P2O5 0.23 0.13

AT22‡ AT24‡ AT25‡ AT27‡ AT29‡ AT33‡ AT45‡ AT50‡ AT53‡

SiO2 53.17 71.86 74.43 49.75 72.99 75.89 48.47 50.63 68.83

TiO2 1.85 0.46 0.48 3 0.43 0.26 1.39 2.29 0.91

Al2O3 16.6 14.99 13.15 14.64 14.45 13.1 18.62 15.48 15.66 FeO 9.75 1.89 2.94 13.02 1.84 1.63 7.91 10.68 3.51 MnO 0.17 0.03 0.05 0.22 0.03 0.02 0.14 0.18 0.05 MgO 5.31 0.42 0.32 4.38 0.39 0.28 6.37 6.41 0.97 CaO 5.87 0.66 0.7 7.72 0.79 0.76 11.54 8.6 1.85

Na2O 5.8 7.13 6.3 4.93 7.15 4.91 2.93 3.95 7.16

K2O 1.01 2.14 1.77 0.71 1.59 2.9 1.11 1.18 1.31

P2O5

AT56‡ AT57‡ AT61‡ AT65‡ AT73‡ AT77‡ AT78‡ AT83‡ AT100‡

SiO2 76.72 67.81 48.5 72.3 50.58 75.02 70.21 56.94 55.55

TiO2 0.22 0.89 1.74 0.3 1.98 0.28 0.62 0.61 1.23

Al2O3 12.65 15.74 16.25 13.8 16.75 13.79 14.14 18.03 15.52 FeO 1.16 3.33 9.29 2.37 9.47 1.82 3.87 6.13 11.15 MnO 0.02 0.05 0.16 0.04 0.15 0.03 0.06 0.12 0.16 MgO 0.22 1 7.02 0.14 7.11 0.24 0.48 4.59 3.94 CaO 0.71 1.26 11.27 0.45 9.42 0.2 1.23 6.67 5.12

Na2O 4.8 7.35 3.6 5.89 4.66 5.85 6.65 3.47 6.59

K2O 3.14 1.97 0.46 3.31 0.28 2.7 1.85 1.56 0.28

P2O5

283

Table D1 (Continued): AT101‡ AT114‡ AT140B‡ AT143‡ AT144‡ AT145‡ AT80-20*

SiO2 52.77 53.28 68.5 76.99 51.52 77.54 51.02

TiO2 1.4 1.08 0.8 0.25 1.91 0.18 0.97

Al2O3 15.61 15.72 16.23 12.65 16.35 12.65 16.97 FeO 12.45 11.61 3.12 1.41 11.7 0.95 11.66 MnO 0.2 0.17 0.05 0.02 0.19 0.02 0.19 MgO 4.53 5.12 0.92 0.16 4.96 0.1 4.81 CaO 8.38 9 2.25 0.49 7.94 1.08 10.51

Na2O 3.45 2.9 6.26 5.54 4.84 4.32 2.65

K2O 0.38 1.12 1.97 2.19 0.32 3.05 1.16

P2O5 0.07 0.10 Total 100.04

AT80-27* AT80-27A* AT80-28* AT80-32§ AT80-43* AT80-43A*

SiO2 59.39 51.74 58.19 58.2 52.05 51.68

TiO2 1.18 1.21 1.10 0.65 0.87 1.14

Al2O3 14.61 15.38 15.79 18.66 17.42 15.29 FeO 11.29 13.93 12.58 5.72 10.16 12.40 MnO 0.33 0.29 0.23 0.15 0.15 0.26 MgO 3.18 6.18 3.32 4.44 5.36 5.41 CaO 4.37 5.43 2.81 6.83 11.23 10.60

Na2O 4.60 3.84 4.48 4.03 2.29 2.52

K2O 1.05 1.88 1.32 1.12 0.47 0.70

P2O5 0.00 0.12 0.18 0.08 0.00 0.00 Total 100.00 100.00 100.00 100.00 100.00

AT80-48* AT80-54* AT80-78* AT80-89* CH9-8† CH9-16† CH9-18†

SiO2 50.99 52.08 52.28 51.14 59.97 49.00 51.24

TiO2 0.73 0.93 1.46 0.92 1.79 1.84 2.26

Al2O3 18.15 16.84 14.08 16.35 14.20 16.40 15.98 FeO 9.93 11.55 15.13 11.13 8.66 9.82 11.33 MnO 0.16 0.19 0.23 0.22 0.17 0.15 0.20 MgO 5.75 4.86 4.38 6.00 3.30 7.40 5.15 CaO 11.28 8.53 8.46 11.34 4.63 11.23 8.52

Na2O 2.40 3.60 2.65 2.51 4.33 3.49 5.07

K2O 0.53 1.33 1.18 0.39 2.30 0.31 0.19

P2O5 0.08 0.09 0.15 0.00 0.26 0.11 0.29 Total 100.00 100.00 100.00 100.00 99.61 99.75 100.23

284

Table D1 (Continued): CP8-16B† CH9-43D§ CH9-52A§ CP98§ CP916C‡ HO9-2D§ HO9-9†

SiO2 56.14 49.3 51.69 53.59 56.53 55.45 53.65

TiO2 0.92 1.85 1.62 0.87 0.9 1.84 1.13

Al2O3 14.59 16.9 18.68 16.06 16.65 16.32 15.45 FeO 10.72 10.28 7.76 11.33 10.15 7.78 12.84 MnO 0.07 0.18 0.15 0.22 0.17 0.12 0.21 MgO 5.35 6.9 5.41 5.44 4.33 3.38 6.74 CaO 7.00 9.5 10.2 5.76 7.03 7.49 4.67

Na2O 3.65 3.92 3.99 4.7 3.27 6.7 5.29

K2O 0.41 0.47 0.23 1.12 0.25 0.54 0.33

P2O5 0.12 0.17 0.24 0.13 Total 98.97 100.44

HO9-14B§ HO922§ HO9-23C§ HO925B§ HO9-42† HO9-49B†

SiO2 49.04 49.79 48.86 49.37 50.15 51.39

TiO2 1.53 1.83 1.03 0.96 1.07 2.52

Al2O3 18.45 15.47 18.84 18.67 13.71 18.91 FeO 8.6 10.03 9.28 8.84 7.71 16.57 MnO 0.18 0.16 0.17 0.13 0.13 0.31 MgO 8.24 8.79 10.06 9.96 6.05 6.86 CaO 10.3 8.04 7.09 6.81 17.40 6.77

Na2O 3.38 3.84 3.57 3.78 2.68 4.45

K2O 0.2 1.27 0.83 0.8 0.25 1.74

P2O5 0.22 0.15 0.21 0 0.10 0.25 Total 99.25 109.77

285

Table D1 (Continued): HO962A§ MB80-24† MP80-2† MP80-8† MP10§ SB1A§ SB80-20§

SiO2 49.47 78.46 53.46 45.91 52.85 49.7 48.98

TiO2 1.64 0.44 0.96 0.77 0.99 1.46 1.47

Al2O3 18 13.36 16.17 14.61 16.79 15.49 18.23 FeO 7.48 2.10 12.01 8.56 12.86 8.94 9.72 MnO 0.14 0.16 0.18 0.43 0.26 0.12 0.34 MgO 5.48 2.33 5.70 6.19 4.79 8.52 6.49 CaO 13.34 0.40 4.73 16.82 6.61 11.47 9.2

Na2O 3.08 3.06 4.70 4.39 5.95 3.82 4.95

K2O 0.79 0.49 2.83 0.48 0.11 0.23 0.22

P2O5 0.11 0.07 0.14 0.11 0.13 0.14 Total 100.87 100.88 98.27 § P2O5 data is from Rubenstone, 1984 and other data is from Yogodzinski et al., 1993; † indicates data is from Rubenstone, 1984; * indicates data is from Shelton, 1986; ‡ from Yogodzinski et al., 1993; additional major element data for Attu can be found in Yogodzinski et al., 1993.

286

Table D2: Attu Island Trace Element Data AT3‡ AT5‡ AT6‡ AT8‡ AT11‡ AT12‡ AT16‡ AT17‡ AT21‡ Sc 9 10.6 49 30.4 45 17.9 37 13.1 26 V Cr 8 9 5 33 137 7 119 3 34 Co 9 7 34 40.9 40 1 43 15.9 24 Ni 7 6 1 48 34 3 69 9 12 Cu Zn Rb Sr 116 543 531 229 147 Y Zr Nb Cs 0.15 0.08 0.15 0.11 0.15 0.02 0.13 0.13 0.03 Ba 279 259 91 197 597 69 82 300 308 La 36.5 22.6 5.11 9.16 8.66 20.48 11.9 19.8 23.2 Ce 93.2 58.3 14 22.47 24.5 52.4 33.9 52.6 50.9 Pr Nd 58.9 35.1 24.7 32.5 Sm 15.4 8.38 4.03 4.46 5.33 9.3 7.69 8.47 9.41 Eu 3.23 1.96 1.82 1.46 1.87 3.09 2.37 3.38 2.86 Gd Tb 2.84 1.52 0.87 0.91 1.11 1.74 1.67 1.53 1.8 Dy Ho Er Tm Yb 12.6 6.85 2.04 3.03 3.46 5.17 5.32 4.92 6.64 Lu 1.92 0.998 0.28 0.444 0.48 0.72 0.75 0.71 0.99 Hf 16.2 9.8 1.82 3.54 3.79 5.56 6.28 6.54 7.86 Ta 3.01 0.97 0.56 0.33 0.54 1.08 0.87 1.02 0.99 Pb Th 5.19 2.29 0.18 0.47 0.43 1.25 0.74 1.24 1.71 U 2.37 1.12 0.12 0.09 0.76 0.23 0.56 0.91

287

Table D2 (Continued): AT22‡ AT24‡ AT25‡ AT27‡ AT29‡ AT33‡ AT45‡ AT50‡ AT53‡ Sc 30 5.3 13.2 38 6 3.1 32 41 9.2 V Cr 38 5 7 8 5 27 241 124 4 Co 36 9.4 20.8 38 21 9.5 36 40 25.4 Ni 37 6 5 14 4 4 63 46 3 Cu Zn Rb Sr 158 479 399 Y Zr Nb Cs 0.36 0.05 0.06 0.02 0.02 0.25 0.25 0.23 0.1 Ba 212 280 262 97 276 445 268 246 301 La 13.7 31.13 36.47 11.4 26.8 20.62 5.73 9.3 18.55 Ce 34.6 78.64 94.23 32 69.1 50.76 15.8 25.1 47.86 Pr Nd 23.2 49.2 57.2 22.3 41.8 21.8 18.2 30.3 Sm 6.11 11.5 13.67 6.83 9.87 4.61 3.39 5.52 7.16 Eu 1.98 2.54 3.09 2.24 2.36 0.88 1.21 1.78 2.5 Gd Tb 1.2 1.77 2.63 1.37 1.8 0.83 0.73 1.22 1.25 Dy Ho Er Tm Yb 3.83 7.05 11.73 4.35 8.2 4.34 2.33 3.93 5.18 Lu 0.54 1.01 1.69 0.6 1.19 0.62 0.32 0.55 0.76 Hf 5.17 9.86 15.65 4.61 11.8 5.2 2.49 4.17 12.59 Ta 0.65 2.19 2.76 0.61 1.82 1.5 0.33 0.52 1.11 Pb Th 1.17 2.45 3.42 0.57 2.47 5.39 0.34 0.45 1.79 U 0.49 1.04 1.62 0.23 1.28 2.23 0.23 0.18 0.92

288

Table D2 (Continued): AT56‡ AT57‡ AT61‡ AT65‡ AT73‡ AT77‡ AT78‡ AT83‡ AT100‡ Sc 2 8.6 38 7 42 7.2 11.4 19 29 V Cr 5 4 348 2 263 7 10 111 13 Co 4 4.5 45 20 44 12.2 4.2 22 29 Ni 2 8 171 1 127 5 4 51 10 Cu Zn Rb Sr 414 355 111 Y Zr Nb Cs 0.15 0.09 0.86 0.17 0.14 0.05 0.08 0.17 0.05 Ba 304 331 125 349 38 356 328 412 31 La 20.4 30.51 5.22 32.7 7.36 30.86 32.25 6.85 3.07 Ce 53 75.08 15.8 79.8 19.6 76.56 81.25 15.8 8.54 Pr Nd 26.4 47.6 12.4 39.1 15.6 46.2 48.9 10.8 8.1 Sm 5.34 11.09 4.16 9.63 5.1 10.42 11.83 2.52 3.02 Eu 0.69 3 1.41 1.54 1.78 2.1 3.09 0.74 0.99 Gd Tb 0.87 1.92 0.95 1.85 1.11 1.93 2.08 0.37 0.79 Dy Ho Er Tm Yb 5.05 8.71 3.2 8.65 4.28 8.94 9.31 1.5 3.03 Lu 0.7 1.27 0.43 1.28 0.68 1.27 1.33 0.215 0.439 Hf 5.94 14.48 3.21 14.19 4 12.53 12.07 2.43 2.16 Ta 2.43 1.84 0.31 2.81 0.46 2.26 2.21 0.12 0.13 Pb Th 5.02 2.35 0.23 4.25 0.16 3.8 2.58 0.83 0.16 U 1.88 1.23 0.09 1.83 0.3 1.78 0.86 0.53 0.06

289

Table D2 (Continued): AT101‡ AT114‡ AT134‡ AT135‡ AT137‡ AT138‡ AT140B‡ AT143‡ Sc 36 42 36.8 33.4 34 28.4 6.2 13 V Cr 13 26 7 31 45 77 7 13 Co 40 37 37.7 34.9 29.5 33.7 15.1 6 Ni 13 14 11 17 18 68 3 2 Cu Zn Rb Sr 231 172 140 249 316 322 Y Zr Nb Cs 0.12 0.13 0.06 0.57 0.22 0.28 0.11 0.02 Ba 84 167 72 224 406 274 239 335 La 3.09 2.4 10.86 7.83 6.73 11.57 18.11 24.6 Ce 8.75 6.36 27.83 19.61 19.05 26.91 41.45 54 Pr Nd 6.3 21.9 14.4 17.7 23.2 26.4 Sm 3.04 2.22 6.45 4.41 4.59 4.95 5.58 5.34 Eu 1.09 0.75 2.07 1.41 1.3 1.43 1.28 0.69 Gd Tb 0.8 0.58 1.35 0.96 1.04 1.09 1 0.87 Dy Ho Er Tm Yb 2.87 2.37 3.99 3.21 3.19 4.06 4.63 5.05 Lu 0.422 0.358 0.581 0.46 0.484 0.607 0.68 0.7 Hf 2.06 1.6 3.93 3.32 3.49 5.24 11.11 5.18 Ta 0.27 0.11 0.55 0.38 0.51 0.6 1.48 1.26 Pb Th 0.19 0.24 0.47 0.52 0.51 0.82 2.75 5.23 U 0.05 0.1 0.25 0.08 0.23 0.21 1.31 2.13

290

Table D2 (Continued): AT144‡ AT145‡ AT80-20* AT80-27A* AT80-27* AT80-28* AT80-32‡† Sc 36 2 42.3 43.3 27.9 29.7 19 V Cr 18 5 37.4 7.9 6.1 3.7 105 Co 28 4 50.4 54.5 11.5 33 21 Ni 7 2 19.6 18.9 0 0 40 Cu Zn Rb 11.3 Sr 286 Y 9.86 Zr 96.5 Nb Cs 0.18 0.37 1.24 0.35 0.11 0.17 0.1 Ba 55 330 251.59 532 219 296 375 La 3.98 25.44 1.92 3.66 6.54 5.82 8.8 Ce 11.2 54.4 5.68 9.44 16.94 18.29 20.1 Pr Nd 9.5 24.1 5.33 6.11 13.5 15.55 12.7 Sm 3.43 5.53 2.07 2.78 4.68 4.21 2.72 Eu 1.41 0.79 0.81 0.95 1.44 1.5 0.78 Gd Tb 0.87 0.97 0.54 0.62 1.03 1.02 0.4 Dy Ho Er Tm Yb 2.98 5.05 2.08 2.48 3.77 3.77 1.56 Lu 0.42 0.73 0.28 0.27 0.61 0.51 0.21 Hf 2.22 5.5 1.19 1.53 3.1 3.03 2.69 Ta 0.3 1.51 5.18 1.96 0.28 2.82 0.2 Pb Th 0.14 5.38 0.27 0.45 0.83 0.72 1.14 U 2.45 0.11 0.24 0.49 0.44 0.71

291

Table D2 (Continued): AT80-43A* AT80-43* AT80-48* AT80-54* AT80-71* AT80-78* Sc 48.79 42.15 42.12 42.06 42.1 47.7 V Cr 71.21 62.29 67.3 35.62 59.5 7.4 Co 19.27 17.73 46.38 51.95 54.5 60.1 Ni 26.77 22.51 20.81 16.62 20.7 16.1 Cu Zn Rb Sr Y Zr Nb Cs 0.27 0.12 0.2 0.67 0.14 0.28 Ba 185 199 124 196 64 293 La 2.24 1.62 1.58 2.12 1.6 2.9 Ce 6.05 4.6 4.49 5.71 5.21 7.84 Pr Nd 5.08 1.99 3.33 3.31 3.43 6.47 Sm 2.23 1.69 1.39 1.99 1.72 2.72 Eu 0.81 0.64 0.62 0.77 0.64 1.02 Gd Tb 0.57 0.44 0.4 0.47 0.46 0.71 Dy Ho Er Tm Yb 2.32 2.02 1.55 2.13 1.87 3.01 Lu 0.37 0.31 0.22 0.26 0.25 0.4 Hf 1.24 0.92 0.81 1.1 1.02 1.58 Ta 0.29 0.29 2.52 3.98 4.31 4.66 Pb Th 0.19 0.13 0.14 0.19 0.18 0.24 U 0.12 0.07 0.12 0.21 0.06 0.16

292

Table D2 (Continued): AT80-89* CH9-8† CH9-16† CH9-18† CH9-43D† CH9-52A† Sc 47.5 24 44.9 39.8 40.2 40.2 V Cr 56.6 13.9 282 107 42.1 114 Co 18.6 22.6 48.1 37.6 44.3 42.2 Ni 15.7 23 86.7 63.2 51.9 59.1 Cu Zn Rb 21.3 4.3 3.01 8.22 3.33 Sr 48.1 275 174 334 208 Y 30.9 36.2 54.2 40.2 30.2 Zr 83.7 107 181 110 106 Nb Cs 0.11 0.09 0.17 0.03 0.2 0.08 Ba 58 540 15 8.7 56 16 La 1.5 4.7 4.37 9.51 5.07 4.55 Ce 4.02 14.8 15.2 27.2 15.2 14.9 Pr Nd 2.25 12.2 13.5 21.9 13 12 Sm 1.63 4.33 4.05 6.73 4.66 3.96 Eu 0.64 1.45 1.52 2.25 1.57 1.46 Gd Tb 0.44 1.07 1.03 1.58 1.18 1.03 Dy Ho Er Tm Yb 1.74 3.37 3.01 4.91 3.68 2.98 Lu 0.26 0.502 0.449 0.707 0.562 0.448 Hf 0.81 2.88 3.49 5.12 3.7 3.1 Ta 0.14 0.28 0.17 0.55 0.25 0.24 Pb Th 0.16 0.328 0.215 0.605 0.404 0.27 U 0.02 0.69 0.302 0.362 0.394 0.24

293

Table D2 (Continued): CP98‡† CP8-16B† CP916C‡ HO9-2D† HO9-9† HO9-14B‡† HO922‡† Sc 40 35.6 39 36.7 53.2 36 37 V Cr 18 40.6 24 236 22.4 363 342 Co 36 37.8 32 34.8 39.3 41 44 Ni 9 25.4 16 100 39.7 187 168 Cu Zn Rb 13.4 6.77 9.84 2.22 4.32 18.8 Sr 255 236 269 225 110 229 267 Y 22.4 27.3 33.4 24.2 28.3 28.3 Zr 51 34.2 153 52.1 90.5 125 Nb Cs 0.06 0.21 0.07 0.12 0.29 0.28 0.3 Ba 392 19 42 39 65 35 204 La 4.51 6.12 6.36 7.97 2.89 6.26 6.1 Ce 11.3 16.5 15.3 24.7 9.38 18.4 18.8 Pr Nd 8.65 11.9 16.3 7.88 13.4 14.9 Sm 2.75 4.34 3.47 4.45 2.66 3.75 4.51 Eu 0.93 1.41 1.09 1.75 0.938 1.36 1.57 Gd Tb 0.57 0.898 0.79 1.11 0.696 0.94 1.04 Dy Ho Er Tm Yb 2 2.54 2.69 3.69 2.83 2.73 3.23 Lu 0.297 0.39 0.359 0.552 0.436 0.422 0.487 Hf 1.58 1.79 1.62 4 2.02 3 3.41 Ta 0.05 0.038 0.04 0.36 0.059 0.37 0.26 Pb Th 0.11 0.316 0.18 0.472 0.496 0.38 0.3 U 0.15 0.337 0.06 0 0.227 0.31 0.77

294

Table D2 (Continued): HO9-23C† HO925B‡† HO9-42† HO9-49B† HO962A‡† MB80-24† Sc 39.5 35 31.7 35.8 32 12 V Cr 153 269 652 200 328 6.41 Co 40.7 43 44.5 41 40 18.6 Ni 104 103 169 109 206 5.06 Cu Zn Rb 3.17 11 5.27 15.3 10.9 5.12 Sr 227 313 183 121 306 41.4 Y 46.9 26.7 22.9 57.3 24.2 343 Zr 142 41.6 54.9 185 129 104 Nb Cs 0.33 0.18 0.18 0.15 0.14 0.05 Ba 7.2 104 12 175 87 8 La 5.93 2.33 3.07 7.49 9.64 4.81 Ce 18.1 6.73 9.32 22.8 22.1 14 Pr Nd 14.1 4.5 7.59 18.9 15 12 Sm 4.82 2.08 2.51 5.85 4.25 3.3 Eu 1.78 0.8 0.957 2.03 1.3 0.824 Gd Tb 1.14 0.5 0.597 1.39 0.81 0.726 Dy Ho Er Tm Yb 3.63 1.69 1.68 4.48 2.75 3.09 Lu 0.565 0.253 0.266 0.679 0.41 0.508 Hf 3.76 1.41 1.95 5 3.34 4.12 Ta 0.21 0.12 0.054 0.36 1.2 0.31 Pb Th 0.353 0.1 0.168 0.223 0.9 0.797 U 0.205 0.189 0.259 0.271 0.43 0.401

295

Table D2 (Continued): MP80-2‡ MP80-8† MP10‡† SB1A‡† SB80-5A‡† SB80-20† Sc 43 33.6 43 36 36 28.2 V Cr 15 802 32 408 4.07 161 Co 38 55.2 43 44 38 33 Ni 14 242 17 185 153 84.3 Cu Zn Rb 5.19 2.13 4.87 21.8 3.79 Sr 494 291 235 183 206 188 Y 12.7 23.1 25.1 43.5 45.2 Zr 62 56.1 79.4 144 144 Nb Cs 0.3 0.11 0.01 0.19 0.36 0.27 Ba 1526 128 16 40 65 23 La 4.89 4.66 6.53 3.9 6.91 7.87 Ce 12.9 13.4 16.3 12.7 20.3 21.1 Pr Nd 9.2 10.3 10.1 14.5 14.5 Sm 2.99 2.52 3.26 3.22 4.37 4.58 Eu 0.95 0.86 1.27 1.19 1.54 1.61 Gd Tb 0.63 0.501 0.64 0.81 1.09 1.17 Dy Ho Er Tm Yb 2.26 1.44 2.22 2.47 3.3 3.85 Lu 0.332 0.249 0.308 0.385 0.518 0.56 Hf 1.83 1.71 1.61 2.47 3.54 4.12 Ta 0.07 0.073 0.06 0.18 0.3 0.26 Pb Th 0.49 0.929 0.44 0.22 0.44 0.47 U 0.11 0.433 0.23 0.12 0.07 0.374 Note: * indicates data is from Shelton, 1986; † indicates data is from Rubenstone, 1984; ‡ indicates data is from Yogodzinski, 1993; and ‡† indicates Rb, Y, and Zr data are from Rubenstone, 1984 and all other data is from Yogodzinski, 1993; all other data is from Yogodzinski, 1993.

296

Table D3: Aleutian Arc Major Element Data Listed By Island ADAK AD-12-1 Shagak 77 Adak G110 BB8-25 BW8-1 BW7-28 BW8-R39B

SiO2 51.62 55.82 53.21 46.95 62.22 59.27 48.84

TiO2 0.82 0.73 0.86 0.91 0.79 1.11 0.84

Al2O3 19.26 19.02 16.82 23.20 17.35 17.00 17.00 FeO 9.59 8.72 9.77 8.22 6.11 6.79 11.26 MnO 0.27 0.25 0.16 0.13 0.12 0.14 0.17 MgO 4.08 5.22 5.93 3.27 2.41 1.76 5.61 CaO 9.20 6.55 8.40 12.59 5.29 4.53 10.58

Na2O 3.45 4.05 2.71 2.92 4.20 4.59 3.32

K2O 1.43 0.59 1.37 0.36 3.03 3.22 0.83

P2O5 0.22 0.29 0.12 0.12 0.20 0.39 0.15 Total 99.95 101.23 99.33 98.65 101.70 98.81 98.61

CC6-63 CC76-74 CM8-1D FB8-2 FB8-19 HB7-6 HB76-10 HB7-10

SiO2 63.73 53.41 56.56 51.48 53.53 54.69 58.42 49.58

TiO2 0.91 0.92 0.88 1.02 0.97 0.82 0.92 0.90

Al2O3 16.87 17.22 13.79 16.20 16.64 18.25 17.24 16.78 FeO 5.71 9.12 10.12 12.15 10.31 8.56 7.80 9.38 MnO 0.13 0.16 0.19 0.21 0.21 0.13 0.14 0.16 MgO 1.42 4.50 6.01 5.44 3.76 5.40 3.72 7.40 CaO 3.04 9.24 7.91 9.86 8.09 9.07 6.79 10.70

Na2O 4.74 3.47 1.89 3.35 4.58 3.31 3.69 2.85

K2O 4.39 1.56 0.76 1.42 1.58 1.35 2.59 0.88

P2O5 0.32 0.25 0.14 0.21 0.32 0.24 0.29 0.17 Total 101.27 99.84 98.25 101.33 99.98 101.81 101.61 98.80

297

Table D3 (Continued): HB7-14A HB7-16 HB6-31A HB76-83J HB6-120 HB76-125 HB6-150 HB-159

SiO2 75.22 52.23 72.59 55.91 57.01 57.07 58.98 53.79

TiO2 0.13 0.74 0.45 0.73 0.83 1.13 0.94 1.08

Al2O3 12.10 16.90 15.16 16.37 17.23 16.67 16.81 17.37 FeO 1.56 8.39 2.55 9.05 7.93 8.40 8.38 9.60 MnO 0.01 0.14 0.06 0.18 0.13 0.13 0.16 0.17 MgO 0.14 5.82 1.16 7.55 4.89 4.83 3.86 4.69 CaO 0.72 9.37 3.77 8.06 7.94 7.98 6.27 8.19

Na2O 3.51 3.02 5.19 3.04 3.34 3.60 4.08 3.53

K2O 4.59 1.12 0.73 0.84 1.94 2.02 2.19 1.00

P2O5 0.03 0.16 0.09 0.31 0.18 0.38 0.34 0.31 Total 98.01 97.91 101.74 102.03 101.41 102.19 102.01 99.74

HB5-166 HB6-174 HB6-181 HB5-193 HB5-207 I8-1 MR8-7b MV80-24

SiO2 54.60 67.41 62.38 60.12 69.52 56.75 51.14 55.80

TiO2 0.91 0.37 0.78 0.90 0.47 0.85 0.65 1.01

Al2O3 17.55 15.57 17.11 16.76 15.96 17.99 19.85 17.25 FeO 7.42 4.01 6.21 6.58 4.31 7.46 9.38 8.66 MnO 0.13 0.05 0.11 0.11 0.09 0.14 0.18 0.15 MgO 3.85 1.24 2.94 3.32 1.91 3.48 3.82 3.74 CaO 6.89 3.57 5.95 6.21 4.09 7.23 8.44 7.85

Na2O 3.92 4.13 3.98 3.64 3.97 3.75 3.64 3.07

K2O 1.56 2.57 2.79 2.95 3.09 1.49 0.76 1.90

P2O5 0.25 0.14 0.20 0.23 0.11 0.28 0.21 0.31 Total 97.06 99.06 102.45 100.82 103.52 99.43 98.07 99.75

298

Table D3 (Continued): KAGALASKA K7-19A K7-22 Kag7-32 Kag7-32A Kag7-43 Kag 34 K5-35b Kg-12-1

SiO2 62.79 56.56 65.03 67.54 50.38 64.60 62.87 49.04

TiO2 0.70 0.99 0.42 0.45 0.88 0.50 0.53 1.00

Al2O3 16.35 17.92 15.95 16.19 19.90 17.13 18.35 18.87 FeO 4.50 8.05 3.74 3.68 9.91 4.58 5.16 12.12 MnO 0.08 0.15 0.07 0.06 0.18 0.10 0.09 0.26 MgO 1.76 2.59 2.00 2.06 5.03 2.39 2.73 4.68 CaO 2.57 6.64 4.60 4.38 8.75 5.02 6.10 10.24

Na2O 3.50 4.19 3.98 3.88 3.71 3.85 4.30 3.06

K2O 6.43 2.34 2.51 2.98 1.21 2.32 1.23 1.08

P2O5 0.13 0.56 0.14 0.12 0.25 0.17 0.14 0.27 Total 98.82 99.99 98.44 101.35 100.20 100.66 101.51 100.61

Kg-12-2 Kg-12-3 Kg-12-4 Kg-12-5 Kg-12-6 Kg-12-7 Kg-12-7b

SiO2 46.97 60.92 53.79 60.74 55.58 51.91 62.02

TiO2 0.94 0.49 0.94 0.62 0.81 1.19 0.95

Al2O3 19.88 18.38 17.56 18.07 17.20 17.39 15.88 FeO 11.84 5.27 10.46 7.02 8.22 9.25 8.89 MnO 0.18 0.10 0.25 0.07 0.30 0.13 0.17 MgO 4.99 1.56 4.76 3.05 4.84 7.12 7.31 CaO 13.28 4.21 7.88 6.66 5.65 8.23 0.72

Na2O 2.47 4.12 3.21 3.44 3.41 3.38 1.01

K2O 0.11 4.66 1.46 0.88 3.46 1.92 2.43

P2O5 0.14 0.21 0.32 0.14 0.18 0.29 0.39 Total 100.79 99.92 100.64 100.68 99.66 100.81 99.77

299

Table D3 (Continued): KAGALASKA KANAGA Kg-12-8 Kg-12-9 Kg-12-10 Kg-12-11 KAN80-2 KAN8-4A

SiO2 48.86 50.75 49.68 58.46 49.80 56.04

TiO2 0.92 0.88 0.73 0.77 0.71 0.68

Al2O3 18.12 15.34 14.78 14.85 16.02 18.65 FeO 12.85 10.25 10.01 8.59 9.20 8.23 MnO 0.21 0.19 0.18 0.10 0.17 0.16 MgO 7.43 7.94 9.75 4.75 11.57 4.23 CaO 5.42 10.62 8.90 10.22 10.28 8.53

Na2O 4.47 3.67 3.23 2.19 2.35 3.54

K2O 1.43 0.97 1.18 0.01 0.55 1.20

P2O5 0.11 0.24 0.18 0.39 0.12 0.18 Total 99.82 100.84 98.62 100.32 100.77 101.46

Kan80-11A KAN80-7-28 SKAN-13-1 SKAN-13-2 SKAN-13-3 SKAN-13-4

SiO2 54.49 62.94 50.79 58.54 58.38 58.56

TiO2 0.73 0.43 1.02 0.72 0.69 0.69

Al2O3 15.08 16.01 16.34 17.42 17.07 16.99 FeO 7.15 4.49 8.73 6.64 6.31 6.25 MnO 0.14 0.11 0.14 0.13 0.11 0.12 MgO 8.70 1.94 7.44 4.75 4.67 3.82 CaO 7.82 4.46 11.24 8.21 7.82 6.86

Na2O 3.27 3.93 3.51 3.52 3.71 3.68

K2O 1.17 1.87 0.79 1.38 1.33 1.31

P2O5 0.15 0.14 0.46 0.19 0.20 0.21 Total 98.70 96.33 100.47 101.50 100.30 98.48

300

Table D3 (Continued): AMLIA AL-13-1 AL-13-2 AL-13-3 AL-13-4 AL-13-5 AL-13-6 AL-13-7

SiO2 54.70 62.64 65.76 68.73 61.71 58.98 52.73

TiO2 0.87 0.67 0.72 0.73 0.88 0.90 0.87

Al2O3 15.09 14.83 14.56 14.25 14.65 15.35 16.35 FeO 8.83 6.54 5.55 5.35 7.89 8.13 9.97 MnO 0.13 0.12 0.11 0.11 0.14 0.13 0.15 MgO 5.47 3.52 1.65 1.30 3.94 3.25 5.85 CaO 8.37 5.18 1.51 1.31 3.35 6.79 7.06

Na2O 3.66 4.47 5.30 5.25 5.00 4.10 4.21

K2O 0.63 1.69 2.99 3.10 2.23 1.03 1.89

P2O5 0.18 0.18 0.25 0.24 0.28 0.28 0.21 Total 97.92 99.84 98.40 100.37 100.08 98.94 99.28

AL-13-8 AL-13-9 AL-13-10 AL-13-11 AL-13-12

SiO2 55.40 52.56 48.20 50.51 71.84

TiO2 1.37 0.71 0.85 0.85 0.67

Al2O3 15.45 15.39 19.44 18.21 13.85 FeO 12.12 8.20 11.08 11.47 4.24 MnO 0.21 0.16 0.17 0.21 0.10 MgO 4.16 7.24 5.57 5.54 0.91 CaO 5.93 11.22 9.68 8.44 1.63

Na2O 4.26 2.97 3.98 4.31 6.08

K2O 1.55 0.22 0.39 0.67 1.36

P2O5 0.33 0.13 0.15 0.12 0.14 Total 100.78 98.79 99.51 100.34 100.83

301

Table D3 (Continued): AMCHITKA AM1 AM4 TAO AM 33 AM 3105 AM 3603 AM 5308 5668 #65

SiO2 59.38 48.07 61.08 48.72 56.65 51.70 49.96

TiO2 1.14 1.19 0.63 0.64 1.09 1.09 1.14

Al2O3 15.48 16.05 17.28 17.10 16.09 14.86 15.02 FeO 8.33 13.08 5.90 9.89 10.58 12.06 12.02 MnO 0.26 0.22 0.09 0.17 0.24 0.14 0.28 MgO 2.33 5.28 3.21 7.57 4.33 4.38 4.98 CaO 3.99 9.10 6.31 11.96 6.13 7.25 8.20

Na2O 6.00 3.74 3.72 2.04 2.49 4.90 3.21

K2O 1.93 1.01 1.76 0.65 0.81 1.61 2.89

P2O5 0.72 0.44 0.13 0.12 0.60 0.35 0.43 Total 99.57 98.17 100.11 98.86 99.00 98.34 98.14

AM #6260 AM-12-1 AM-12-2 51-L-8 52-Sn-11 51-N-23 51-P-27 51-P-31

SiO2 47.34 55.41 56.40 65.78 60.58 49.07 63.39 60.38

TiO2 1.03 1.07 1.09 0.48 0.57 0.82 0.59 0.67

Al2O3 16.34 15.53 15.53 17.51 18.54 20.60 16.04 18.31 FeO 13.43 10.72 10.71 4.42 5.56 10.98 5.17 6.08 MnO 0.22 0.22 0.25 0.07 0.11 0.17 0.13 0.12 MgO 6.73 3.35 2.55 1.52 2.24 4.47 1.14 3.32 CaO 7.26 4.07 5.48 4.23 7.23 11.87 3.73 7.39

Na2O 5.10 4.01 4.61 2.95 3.53 2.49 4.04 3.60

K2O 0.93 3.41 2.46 2.06 1.21 0.45 4.37 1.71

P2O5 0.22 0.61 0.62 0.08 0.19 0.07 0.27 0.21 Total 98.60 98.39 99.70 99.09 99.76 100.97 98.88 101.80

302

Table D3 (Continued): AMATIGNAK AT-12-2 AT-12-3 AT-12-4 AT-12-5 AT-12-6 AMT-13-10 AMT-13-11

SiO2 49.28 61.14 50.70 57.19 58.26 54.55 52.36

TiO2 1.11 1.18 1.44 0.88 0.89 1.18 1.10

Al2O3 18.97 14.28 15.63 16.20 16.20 16.56 18.65 FeO 9.78 8.40 12.94 8.38 8.03 9.33 9.00 MnO 0.15 0.17 0.21 0.17 0.16 0.16 0.15 MgO 7.23 2.28 4.54 4.07 3.43 2.79 2.94 CaO 10.57 6.53 8.15 6.79 7.03 6.65 7.46

Na2O 2.87 4.68 3.91 4.44 3.89 4.69 4.64

K2O 0.39 1.07 1.64 2.14 1.98 1.81 1.95

P2O5 0.29 0.48 0.49 0.24 0.26 0.46 0.36 Total 100.64 100.21 99.66 100.50 100.12 98.19 98.61

AMT-13-7 AMT-13-8 AMT-13-9A AMT-13-9B AMT-13-9C 52-Sn-141

SiO2 52.98 57.21 56.27 53.49 49.71 56.18

TiO2 1.22 0.88 1.01 0.98 0.89 0.86

Al2O3 14.65 16.16 15.11 16.86 16.71 16.28 FeO 11.40 8.44 12.15 10.71 10.88 8.14 MnO 0.26 0.16 0.21 0.18 0.20 0.16 MgO 4.91 3.88 5.02 5.28 6.38 3.80 CaO 6.55 7.40 6.10 6.95 7.73 5.88

Na2O 5.22 3.87 3.27 3.85 4.50 4.30

K2O 1.48 1.81 1.49 2.30 0.85 2.20

P2O5 0.34 0.25 0.27 0.28 0.17 0.26 Total 99.00 100.06 100.91 100.90 98.01 98.06

303

Table D3 (Continued): AMATIGNAK ULAK 52-Sn-147 52-Sn-160 UL-12-1 UL-12-2 UL-12-3 UL-12-4 UL-12-5

SiO2 53.52 67.53 56.48 50.64 48.05 48.74 54.41

TiO2 1.17 0.36 0.41 0.79 0.91 0.68 1.05

Al2O3 17.59 15.84 12.62 16.73 16.80 18.81 15.14 FeO 10.72 3.42 8.76 11.26 11.36 9.96 10.18 MnO 0.21 0.06 0.12 0.16 0.17 0.13 0.17 MgO 4.31 1.32 7.03 7.11 8.90 5.85 4.57 CaO 7.62 3.55 12.46 8.14 10.56 11.36 7.39

Na2O 4.26 3.98 1.68 4.79 2.75 2.84 4.62

K2O 1.48 2.54 0.24 0.28 0.22 0.64 0.34

P2O5 0.38 0.11 0.20 0.15 0.13 0.18 0.30 Total 101.24 98.70 99.98 100.04 99.86 99.19 98.15

UL-13-6A UL-13-6B UL-13-7 UL-13-8A UL-13-8B 52-Bi-6 52-Bi-12

SiO2 55.84 48.13 63.86 52.34 56.53 47.32 51.71

TiO2 1.10 0.81 0.72 1.53 0.63 0.88 0.92

Al2O3 16.12 16.12 14.40 16.01 13.25 18.10 21.03 FeO 8.87 10.40 7.53 12.38 6.33 12.51 9.09 MnO 0.18 0.15 0.14 0.21 0.19 0.18 0.18 MgO 4.33 8.24 4.05 4.29 6.38 6.51 3.35 CaO 7.58 10.49 5.25 6.77 11.05 8.82 10.72

Na2O 3.76 3.14 3.88 5.11 2.08 2.51 3.44

K2O 1.17 0.70 0.30 1.50 4.71 1.79 0.48

P2O5 0.30 0.13 0.22 0.45 0.19 0.22 0.20 Total 99.26 98.31 100.35 100.60 101.35 98.84 101.12

304

Table D3 (Continued): GAREOLI GAR-13-1 GAR-13-2 GAR-13-3

SiO2 53.94 52.20 53.48

TiO2 0.88 1.04 0.87

Al2O3 20.65 17.74 19.64 FeO 7.80 9.36 7.62 MnO 0.14 0.17 0.15 MgO 2.62 4.66 2.56 CaO 8.72 8.97 8.13

Na2O 4.14 3.71 4.17

K2O 2.17 1.85 2.14

P2O5 0.40 0.37 0.36 Total 101.48 100.07 99.11

KAVALGA KAV-13-1 KAV-13-2 KAV-13-2B KAV-13-3

SiO2 56.69 51.13 52.22 55.86

TiO2 0.80 1.04 1.04 0.91

Al2O3 18.11 19.92 20.22 17.37 FeO 7.61 9.09 9.26 7.39 MnO 0.15 0.21 0.20 0.14 MgO 3.71 3.12 3.14 3.68 CaO 8.07 8.88 9.13 7.77

Na2O 3.64 4.00 4.05 3.79

K2O 1.25 0.91 0.96 1.59

P2O5 0.22 0.18 0.18 0.18 Total 100.24 98.47 100.40 98.68

305

Table D3 (Continued): SKAGUL SKA-13-1 SKA-13-2 SKA-13-2K SKA-13-3 SKA-13-3K SKA-13-4 SKA-13-4B

SiO2 59.43 49.81 50.17 55.79 55.89 53.69 49.70

TiO2 0.59 0.92 0.87 0.59 0.62 0.63 0.93

Al2O3 17.56 19.53 20.10 19.34 19.57 20.39 19.56 FeO 6.04 10.80 10.35 7.65 8.47 9.24 10.67 MnO 0.13 0.17 0.16 0.18 0.23 0.18 0.17 MgO 2.38 5.99 6.39 2.41 2.91 3.48 6.45 CaO 6.37 9.82 9.80 8.11 8.93 9.19 9.48

Na2O 4.02 3.29 3.21 3.86 3.76 3.53 3.14

K2O 1.35 0.58 0.53 0.77 0.72 0.72 0.58

P2O5 0.16 0.15 0.14 0.30 0.31 0.24 0.14 Total 98.03 101.06 101.73 99.00 101.42 101.29 100.82

OGLIUGA GREAT SITKIN SEDANKA OGL-13-1 OGL-13-2 GS80-R13 SED 36D SED 4B

SiO2 48.89 47.63 48.23 54.28 50.49

TiO2 0.97 1.08 0.79 0.92 0.77

Al2O3 19.85 18.17 17.03 15.67 19.76 FeO 10.53 11.41 9.74 8.84 8.65 MnO 0.16 0.20 0.17 0.09 0.12 MgO 5.17 7.94 10.50 5.50 5.18 CaO 9.90 11.79 8.14 8.77 10.38

Na2O 3.23 2.64 2.72 2.99 3.30

K2O 0.62 0.59 1.35 0.65 0.40

P2O5 0.18 0.17 0.09 0.25 0.22 Total 99.51 101.63 98.76 97.96 99.27 Note: All major element data was obtained using Cornell University’s ICP -OES following the method described in appendix A.

306

Table D4: Aleutian Arc Trace Element Data Listed By Island ULAK UL-12-4* UL-13-6B UL-13-8A UL-13-8B 52-Bi-6 52-Bi-12 Sc 28.134 44.033 32.774 19.881 30.570 22.274 V 310.435 314.668 392.816 270.011 344.036 283.298 Cr 77.602 207.478 3.679 362.345 79.192 10.721 Co 35.889 38.464 27.352 24.917 41.294 23.565 Ni 35.102 68.726 7.822 103.003 40.570 10.423 Cu 132.746 112.422 348.492 78.063 159.041 320.394 Zn 73.435 71.139 109.265 46.568 97.176 92.693 Rb 4.761 13.929 28.100 85.693 38.705 7.798 Sr 450.498 490.395 481.904 274.877 709.414 573.178 Y 15.956 15.230 38.185 12.702 18.911 24.127 Zr 58.329 54.051 133.663 67.142 62.279 55.372 Nb 1.288 1.006 4.900 2.060 1.069 1.018 Ba 286.998 299.694 290.642 568.803 242.005 138.901 La 5.021 6.851 12.583 4.583 7.629 6.573 Ce 13.348 15.527 31.991 10.976 18.398 15.959 Pr 2.077 2.415 4.996 1.633 2.953 2.610 Nd 9.182 10.964 22.898 7.476 13.913 12.826 Sm 2.738 2.732 5.771 1.947 3.543 3.484 Eu 0.775 0.915 1.674 0.709 1.181 1.221 Gd 3.032 2.880 5.971 2.220 3.606 3.785 Tb 0.463 0.449 0.958 0.353 0.564 0.624 Dy 2.665 2.583 5.660 2.064 3.215 3.774 Ho 0.521 0.539 1.183 0.439 0.671 0.809 Er 1.537 1.476 3.251 1.215 1.851 2.243 Tm 0.236 0.220 0.498 0.184 0.279 0.340 Yb 1.521 1.340 3.003 1.133 1.716 2.090 Lu 0.175 0.206 0.465 0.178 0.267 0.328 Hf 1.409 1.549 3.596 1.630 1.785 1.684 Ta 0.128 0.068 0.326 0.145 0.070 0.062 Pb 1.755 2.637 3.475 3.515 4.633 3.274 Th 0.704 1.500 1.812 1.058 1.035 0.726 U 0.414 0.654 0.837 0.653 0.508 0.391

307

Table D4 (Continued): AMATIGNAK AT-12-4* AT-12-5* AT-12-6* AMT-13-7 AMT-13-8 AMT-13-10 AMT-13-11 Sc 30.750 20.083 19.174 31.160 23.992 20.553 19.917 V 417.380 216.989 214.458 355.012 212.953 214.346 246.313 Cr 73.720 90.617 81.245 14.121 37.507 6.446 8.977 Co 34.608 21.865 19.489 28.013 21.499 19.785 20.254 Ni 24.341 23.084 18.988 18.091 17.819 7.982 18.520 Cu 290.425 63.419 89.894 253.561 58.989 22.837 198.411 Zn 118.008 70.709 74.720 124.697 74.821 80.650 80.214 Rb 20.480 28.724 30.454 30.596 37.866 51.219 39.475 Sr 547.577 536.466 387.724 510.211 420.668 491.277 668.503 Y 29.417 22.859 26.156 25.375 27.659 39.456 29.818 Zr 119.767 102.893 49.580 89.814 36.876 160.099 117.494 Nb 4.063 5.403 6.315 2.807 3.911 5.899 4.306 Ba 470.630 361.581 410.763 864.808 365.319 424.828 603.691 La 12.001 9.006 10.962 8.760 11.117 13.839 10.974 Ce 33.693 23.507 27.852 22.878 26.038 35.836 26.919 Pr 4.892 3.402 3.826 3.694 3.982 5.439 4.237 Nd 21.550 14.625 16.093 17.353 17.721 24.808 19.264 Sm 6.166 4.122 4.449 4.419 4.343 6.152 4.817 Eu 1.598 1.020 1.154 1.451 1.278 1.821 1.581 Gd 5.410 4.034 4.454 4.627 4.569 6.262 5.042 Tb 0.897 0.651 0.711 0.728 0.740 0.993 0.795 Dy 5.643 4.085 4.318 4.255 4.352 5.842 4.657 Ho 1.152 0.826 0.859 0.882 0.913 1.219 0.975 Er 3.219 2.384 2.460 2.418 2.498 3.366 2.696 Tm 0.495 0.366 0.376 0.367 0.377 0.514 0.412 Yb 3.128 2.347 2.397 2.230 2.281 3.146 2.494 Lu 0.413 0.314 0.256 0.342 0.346 0.489 0.389 Hf 3.354 2.835 1.307 2.398 1.270 3.821 3.041 Ta 0.280 0.351 0.456 0.186 0.264 0.371 0.275 Pb 2.802 4.115 5.181 1.710 3.801 1.526 1.911 Th 1.990 1.461 2.127 1.142 1.841 1.951 1.459 U 0.964 0.880 0.626 0.525 0.602 0.873 0.679

308

Table D4 (Continued): OGLIUGA SKAGUL OGL-13-1 OGL-13-2 SKA-13-4 Sc 21.574 39.802 10.444 V 305.697 378.169 157.587 Cr 12.911 106.053 4.781 Co 31.176 43.173 22.595 Ni 19.737 62.870 7.316 Cu 63.876 135.649 73.835 Zn 75.634 81.084 79.728 Rb 8.618 8.462 12.786 Sr 530.007 596.431 669.856 Y 19.284 17.632 16.538 Zr 80.181 65.257 91.327 Nb 1.383 1.024 1.531 Ba 207.627 208.562 270.518 La 7.323 7.997 8.924 Ce 18.674 20.014 21.732 Pr 2.993 3.185 3.341 Nd 14.125 14.888 14.644 Sm 3.586 3.754 3.348 Eu 1.200 1.250 1.198 Gd 3.563 3.762 3.335 Tb 0.566 0.564 0.496 Dy 3.206 3.085 2.717 Ho 0.669 0.614 0.556 Er 1.807 1.649 1.574 Tm 0.272 0.235 0.240 Yb 1.644 1.426 1.503 Lu 0.253 0.213 0.238 Hf 2.156 1.804 2.294 Ta 0.088 0.071 0.092 Pb 1.410 1.926 1.622 Th 0.729 0.783 0.911 U 0.352 0.330 0.454 Note: * indicates data obtained using Cornell University’s ICP-MS all other data was obtained using Colgate University’s ICP-MS. Methods for both are described in Appendix A.

309

Table D5: Isotope Data Listed By Island 87Sr/86Sr 2σ 143Nd/144Nd 2σ Epsilon Nd 2σ

ADAK AD-12-2 0.703016 ±8 0.513040 ±16 7.83 ±0.30 Shagak 77 0.703834 ±10 0.513012 ±13 7.29 ±0.26 AdakG110 0.703181 ±7 0.513088 ±14 8.77 ±0.28 BB8-25 0.703093 ±7 0.513114 ±11 9.28 ±0.22 BW8-1 0.703271 ±6 0.513063 ±14 8.30 ±0.28 BW8-R27 0.703067 ±10 0.513073 ±14 8.48 ±0.28 BW8-38 0.703231 ±7 0.513045 ±12 7.93 ±0.24 BW8-R39B 0.702986 ±7 0.513097 ±13 8.95 ±0.26 CC-76-74 0.703057 ±6 0.513047 ±16 7.97 ±0.30 CC-76-122 0.703011 ±7 0.513091 ±16 8.84 ±0.30 CM8-1A 0.703448 ±8 0.513040 ±14 7.83 ±0.28 CM8-1D 0.703493 ±14 0.512952 ±18 6.13 ±0.36 CM8-3 0.703360 ±8 0.513034 ±16 7.72 ±0.30 CP80-9 0.703435 ±6 0.5130391 ±14 7.82 ±0.28 FB8-2 0.703087 ±7 0.513079 ±11 8.59 ±0.22 FB8-19 0.703012 ±8 0.513074 ±12 8.51 ±0.24 HB7-6 0.702900 ±8 0.513037 ±13 7.79 ±0.26 HB76-10 0.703145 ±7 0.513060 ±12 8.22 ±0.24 HB7-10 0.703006 ±7 0.513057 ±12 8.16 ±0.24 HB-76-18A 0.703299 ±11 0.513076 ±10 8.54 ±0.20 HB7-19 0.703194 ±7 0.513058 ±12 8.19 ±0.24 HB-76-51 0.703148 ±10 0.513061 ±11 8.26 ±0.22 HB-76-83J 0.703301 ±7 0.513051 ±17 8.06 ±0.34 HB6-120 0.703055 ±7 0.513042 ±16 7.89 ±0.32 HB76-125 0.703134 ±7 0.513067 ±13 8.37 ±0.26 HB-159 0.703124 ±11 0.513050 ±10 8.03 ±0.20 HB6-174 0.703257 ±7 0.512997 ±18 7.00 ±0.36 HB5-193 0.703278 ±14 0.513010 ±13 7.25 ±0.26 HB6-202 0.703149 ±7 0.513096 ±14 8.94 ±0.28 HB5-207 0.513039 ±16 7.83 ±0.30 HB6-260I 0.703029 ±6 0.513085 ±11 8.71 ±0.22 I8-1 0.703176 ±8 0.513038 ±18 7.80 ±0.36 MR8-7B 0.703388 ±6 0.513023 ±13 7.50 ±0.26 MR80-12 0.70317 ±10 0.5130782 ±13 8.59 ±0.26 MV-80-24 0.703095 ±8 0.513086 ±11 8.73 ±0.22 GREAT SITKIN GS80-R13 0.703673 ±6 0.513081 ±13 8.65 ±0.26

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Table D5 (Continued): 87Sr/86Sr 2σ 143Nd/144Nd 2σ Epsilon Nd 2σ KAGALASKA Kag7-32 0.703079 ±11 0.513072 ±17 8.46 ±0.34 Kag7-32A 0.703029 ±13 0.513074 ±16 8.50 ±0.30 Kag 34 0.703072 ±8 0.513078 ±16 8.59 ±0.32 Kag7-50 0.703076 ±10 0.513058 ±11 8.19 ±0.22 KG-12-3 0.703263 ±7 0.513043 ±16 7.90 ±0.30 KG-12-5 0.702951 ±6 0.513091 ±13 8.84 ±0.26 KANAGA Kan80-2 0.703163 ±10 0.513056 ±17 8.16 ±0.34 KAN80-4A 0.703202 ±10 0.513062 ±14 8.27 ±0.28 KAN80-11A 0.703251 ±6 0.513040 ±18 7.83 ±0.36 KAN80-7-28 0.703227 ±7 0.512964 ±16 6.37 ±0.30 SKAN-13-1 0.703019 ±7 0.513051 ±10 8.06 ±0.20 SKAN-13-4 0.703037 ±8 0.513029 ±12 7.63 ±0.24 AMLIA AL-13-2 0.703489 ±10 0.513052 ±17 8.07 ±0.34 AL-13-4 0.704483 ±10 0.513024 ±11 7.53 ±0.22 AL-13-7 0.703665 ±11 AL-13-9 0.703305 ±11 0.513069 ±10 8.40 ±0.20 AL-13-10 0.703536 ±13 0.513099 ±16 8.98 ±0.32 AMCHITKA 51-L-8 0.703277 ±6 0.513074 ±14 8.51 ±0.28 52-Sn-11 0.703369 ±10 0.513040 ±20 7.84 ±0.38 51-N-23 0.703177 ±7 0.513085 ±22 8.72 ±0.42 51-P-27 0.703683 ±11 0.513068 ±17 8.39 ±0.34 51-P-31 0.703075 ±7 0.513084 ±16 8.70 ±0.30 AM-12-2 0.703214 ±6 0.513016 ±13 7.36 ±0.26 AM1 0.703464 ±7 0.512943 ±16 5.95 ±0.32 AM4 TAO 0.703791 ±7 0.513067 ±13 8.36 ±0.26 AM33 0.703385 ±10 0.513050 ±14 8.03 ±0.28 AM 3105 0.703107 ±6 0.513055 ±14 8.14 ±0.28 AM 3603 0.703422 ±6 0.513076 ±12 8.54 ±0.24 AM 5308 0.703438 ±8 0.513014 ±12 7.33 ±0.24 5668 #65 0.703314 ±13 0.513065 ±14 8.33 ±0.28 AM #6260 0.703390 ±7 0.513008 ±18 7.21 ±0.36 SEDANKA SED AB/4B 0.703308 ±8 0.513032 ±12 7.68 ±0.24 SED 36D 0.703340 ±8 0.513003 ±11 7.11 ±0.22

311

Table D5 (Continued): 87Sr/86Sr 2σ 143Nd/144Nd 2σ Epsilon Nd 2σ AMATIGNAK 52-Sn-141 0.703823 ±8 0.513107 ±10 9.15 ±0.20 52-Sn-147 0.703870 ±7 0.513068 ±16 8.38 ±0.30 52-Sn-160 0.703257 ±10 0.513011 ±13 7.28 ±0.26 AT-12-2 0.703129 ±8 0.513102 ±20 9.05 ±0.38 AT-12-3 0.703500 ±13 0.513093 ±13 8.88 ±0.26 AT-12-4 0.703868 ±6 0.513045 ±13 7.94 ±0.26 AT-12-5 0.703962 ±8 0.513086 ±16 8.74 ±0.30 AT-12-6 0.703304 ±10 0.513096 ±12 8.93 ±0.24 AMT-13-7 0.703412 ±10 0.513025 ±16 7.54 ±0.30 AMT-13-8 0.703329 ±8 0.513082 ±7 8.66 ±0.14 AMT-13-10 0.703767 ±7 0.513071 ±10 8.45 ±0.20 AMT-13-11 0.703597 ±11 0.513020 ±16 7.46 ±0.32 ULAK 52-Bi-6 0.703533 ±6 0.513098 ±14 8.97 ±0.28 52-Bi-12 0.703268 ±8 0.513083 ±17 8.68 ±0.34 UL-12-1 0.703889 ±6 0.513094 ±16 8.90 ±0.30 UL-12-2 0.704026 ±10 0.513064 ±13 8.31 ±0.26 UL-12-3 0.703591 ±10 0.513073 ±16 8.49 ±0.32 UL-12-4 0.703198 ±8 0.513044 ±21 7.92 ±0.40 UL-13-6A 0.703256 ±8 0.513051 ±16 8.06 ±0.30 UL-13-6B 0.703372 ±8 0.513031 ±16 7.67 ±0.30 UL-13-7 0.703810 ±7 0.513078 ±12 8.57 ±0.24 UL-13-8A 0.703403 ±7 0.513071 ±13 8.45 ±0.26 UL-13-8B 0.703766 ±10 0.513042 ±16 7.88 ±0.32 GAREOLI GAR-13-3 0.703162 ±10 0.512948 ±17 6.05 ±0.34 KAVALGA KAV-13-1 0.703061 ±7 0.513050 ±14 8.03 ±0.28 KAV-13-2 0.703044 ±7 0.513041 ±13 7.86 ±0.26 KAV-13-3 0.703060 ±8 0.513069 ±9 8.41 ±0.18 SKAGUL SKA-13-1 0.703044 ±7 0.513087 ±14 8.77 ±0.28 SKA-13-2K 0.703036 ±8 0.513082 ±16 8.65 ±0.30 SKA-13-3K 0.703105 ±10 0.513057 ±13 8.17 ±0.26 SKA-13-4 0.703100 ±6 0.513063 ±10 8.30 ±0.20 OGLIUGA OGL-13-1 0.703096 ±8 0.513063 ±10 8.29 ±0.20 OGL-13-2 0.703082 ±7 0.513042 ±14 7.87 ±0.28 Note: Data was obtained using Cornell University’s TIMS; method is described in Appendix A.

312

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