1 Timing of Tectonic Uplift Rate Change at Araki, Vanuatu, Derived from 230

Timing of tectonic uplift rate change at Araki, Vanuatu, derived from 230-Th dating of fossil corals

A Thesis SUBMITTED TO THE FACULTY OF UNIVERSITY OF MINNESOTA BY

Claire Rabine

IN PARTIAL FULFULLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE

Dr. Christina Gallup, Advisor

January 2019

Acknowledgements

This project was greatly improved by the knowledge and comments from my advisor,

Christina Gallup, and my committee, Byron Steinman, Kathryn Schreiner, John Goodge, and Fred Taylor (University of Texas – Austin). Special thanks are owed to Larry

Edwards for use of his laboratory, as well as Xianlei Li and Pu Zhang for their invaluable help (University of Minnesota, Twin Cities). Training and insight about x-ray diffraction by Tsutomu “Shimo” Shimotori is also greatly appreciated (UMD). Most of all, thanks to my fiancé, family, and friends for their continual support throughout this project.

Abstract

New data from precisely leveled and 230Th-dated coral fossils allow for better constraints on the recent (~125 ka) tectonic history of Araki, Vanuatu. Subduction of two large underwater landforms have caused significant deformation to the intraoceanic arc on which Araki is centered and induced a variable and complex tectonic history. Due to the relatively few intraoceanic arcs undergoing similar collisions, an effort to better constrain and understand the tectonic history of Araki is valuable. This study presents an uplift model that builds on previous constraints of the tectonic history of Araki using fossil coral elevation and age data and a reference sea level curve. Data from fifty-nine samples was collected and used to provide chronological constraints for paleo shorelines that in turn offered insight into the tectonic history. The model suggests subsidence at a rate of 3.9 mm/yr from

125 ka to 106 ka, a shift to uplift at 106 ka continuing at a rate of 1.4 to 1.6 mm/yr until 25 – 30 ka, and then increasing to a rate of 4.6 mm/yr until the present. The abrupt changes in vertical tectonics implied by this modeling offers insight into the rapid tectonic variability and possible mechanisms controlling convergent margin tectonics.

Table of Contents Acknowledgements...... i Abstract...... ii List of Figures ...... iv 1. Introduction ...... 1 1.1 Tectonic history of New Hebrides ...... 3 1.2 Determination of past sea levels using coral terraces ...... 8 1.3 History of fossil coral studies on Araki...... 12 1.4 230Th Dating ...... 13 1.5 Coral Sample Selection ...... 17 2. Methods ...... 17 2.1 Surveying and Collection of Samples...... 17 2.2 Process of U-Th Dating ...... 18 2.3 X-Ray Diffraction...... 21 3. Results ...... 25 4. Discussion ...... 27 4.1 Early subsidence from 125 to 106 ka ...... 28 4.2 Long-term uplift rate ...... 29 4.3 Uplift rate since ~20 ka...... 29 4.4 Modeling uplift rate from 20-105 ka ...... 31 4.5 Errors encountered and future work...... 39 5. Conclusions: ...... 44 References: ...... 46 Appendix A: 230Th Dataset ...... 51 Appendix B: XRD Results ...... 54

iii

List of Figures

Figure 1: Vanuatu overview map ...... 4

Figure 2: Espiritu Santo subduction cross-section ...... 5

Figure 3: Papua New Guinea map ...... 9

Figure 4: Sea level curve by Lambeck et al. (2002) ...... 12

Figure 5: Uranium decay chain ...... 14

Figure 6: Calcite and aragonite diffractograms ...... 23

Figure 7: Calcite/aragonite calibration curve ...... 25

Figure 8: Drilling locations near Araki ...... 29

Figure 9: Simplified uplift history ...... 31

Figure 10: Example of model using Sample AF ...... 34

Figure 11: Vertical uplift rates modeled ...... 37

Figure 12: Sea level estimates from model ...... 38

Figure 13: Error due to growth depth of corals ...... 39

Figure 14: Error in uplift rates due to 106 ka data point ...... 41

Figure 15: Error in sea levels due to 106 ka data point ...... 42

Figure 16: Porites distribution histogram ...... 43

1. Introduction

Araki is a small island within a much larger chain of islands in the Pacific known as the Republic of Vanuatu. Vanuatu sits at the convergence of the Pacific and Australian plates (Figure 1), and is known to have an uncommonly variable tectonic history (Collot et al., 1985). The subduction of two massive underwater features, an extinct subduction zone known as the d’Entrecasteaux Zone (DEZ) and an underwater plateau known as the

West Torres Massif (WTM), has had tremendous impact on the structure and tectonic development of the arc. As these features collide and subduct, the islands of Vanuatu uplift or subside in response. Other notable effects include the slowing of plate convergence rates closest to the DEZ and WTM, rapid thickening and uplift of the forearc, and a very shallow trench (Meffre and Crawford, 2001) (see Figure 2).

Additionally, the forearc islands are unusually close to the trench, which means the seismogenic interplate thrust zone is directly beneath Araki at a depth of only 20 – 25 km

(Baillard et al., 2015). It is unusual and fortuitous to have islands positioned to record vertical tectonic deformations directly over this zone, which is greatly influenced by tectonic deformation. It is possible to document this movement using coral fossils from

Araki and other nearby islands, as many species of corals grow at or near sea level and are able to be precisely dated using uranium-thorium radiometric dating (Edwards, 1988).

It has been previously estimated that the highest rate of uplift in Vanuatu is along the western coast of its largest island, Espiritu Santo, just north of Araki, at rates as high as 6 mm/yr (Taylor, 1992). High uplift rates (~ 4 – 6 mm/yr) have been occurring since 1 approximately the Last Glacial Maximum (~20 ka), which preceded slower rates estimated to be about 3.2 mm/yr (Taylor, 1992). There is evidence of subsidence as well, prior to about 170 to 100 ka (Taylor, 1992; Taylor, 1995). However, the timing and magnitude of these changes in rate are unclear. A collection of 59 coral fossils between the ages of 30 and 140 ka collected by Fred Taylor (UT Austin) and Nick Freiburger

(UMD) in 2005 may provide additional constraints on the tectonic history. Eleven samples were dated by Freiburger using 230Th dating as part of his MS thesis (2006), another twelve samples were dated by Tyler Carlson (2008, unpublished), and the remaining thirty-six are dated in this study. They range in age from 32.16 ± 0.04 ka to

141.3 ± 0.3 ka and range in elevation from 32 ± 1 to 183 ± 1 meters above sea level (see

Appendix A for full data set). I present a model that uses ages and elevations of these new coral fossils and constraints from previous work to calculate potential uplift rates and a corresponding sea level history. This history is compared to an existing sea level curve (Lambeck et al., 2002), and the predicted uplift rates are adjusted until a best fit is found. The results of this model suggest that Araki subsided from 125 ka to 106 ka at a rate of about 3.9 mm/yr, followed by a shift to uplift, which continued until about 30 – 25 ka at a rate of 1.4 to 1.6 mm/yr. From 25 – 30 ka to present the uplift rate increased to about 4.6 mm/yr. This better-constrained tectonic history is significant because it offers insights into the rapid tectonic variability, which may be used to better understand mechanisms controlling convergent margin tectonics at other locations around the globe.

1.1 Tectonic history of New Hebrides

The Republic of Vanuatu is a Pacific island nation located approximately 2,500 kilometers east of northern Australia on the Pacific plate. It is an archipelago consisting of roughly 80 islands of predominantly volcanic origin, spanning 1,400 km along the convergent boundary between the Pacific and Australian plates. The Vanuatu island archipelago is also referred to as the New Hebrides island arc system (see Figure 1 for an overview of the area, as well as for the names of individual islands, plates, and features), which sits on the Pacific plate margin overlying the subducting Australian plate. The convergence rate of the Australian plate subducting beneath the Pacific plate varies significantly, ranging from 9 – 12 cm/yr to 2.5 – 4 cm/yr (Baillard et al., 2015). The area of slowest convergence is in the central section of the arc where the subduction of two sizeable bathymetric features, the d’Entrecasteaux Zone (DEZ) and the West Torres

Massif (WTM), is occurring. The DEZ is an extinct, north-facing subduction zone that includes a twin-spined ridge composed of a series of ancient stratovolcanoes on the southern seamount chain and lavas on the northern ridge. It is estimated that it began colliding with the New Hebrides arc at 2-3 Ma (Meffre and Crawford, 2001; Green and

Collot, 1994). The WTM is a massive submarine plateau of largely unknown origin, rising up to 4 km above the surrounding sea floor and covering an area of more than

35,000 km2. It likely began colliding about 1 Ma (Meffre and Crawford, 2001).

Figure 1: Map of Central New Hebrides island arc, showing major islands and bathymetry (Adapted from Taylor, 1992).

4 Subduction of these features has significant effects on the structure of the arc.

Other island arcs undergoing large collisions have similar structures as well (Meffre &

Crawford, 2001). The most obvious is that in the central part of the arc, between latitudes

13°S and 17°S, the conventional single island chain configuration is instead composed of three parallel chains: uplifted outer forearc islands to the west (Espiritu Santo and

Malakula), active volcanic arc islands in the center (Aoba and Ambrym), and uplifted back-arc islands dominated by reefs to the east (Maewo and Pentecost) (Collot et al.,

1985). Additionally, Espiritu Santo and Malakula are unusually close to the convergent plate boundary, 30 km from the trench, a location normally corresponding to the middle of the trench-arc slope in most intraoceanic island arcs (Meffre and Crawford, 2001).

Subduction of the DEZ and WTM also elevates the island arc system, causing Espiritu

Santo and Malakula to be highly uplifted and the trench to the west of these islands to be abnormally shallow, nearly nonexistent for about 500 km near 15°S (Meffre and

Crawford, 2001) (Figure 2). The formation and depth of Aoba Basin, in which the

Figure 2: Simplified cross section of southern Espiritu Santo, showing collision of DEZ. Dashed line represents typical non-collisional profile; shaded areas show uplifted parts of the arc. Vertical exaggeration is ~6.5. Adapted from Meffre and Crawford (2001).

volcanic island Aoba is centered, is also attributed to the subduction of the DEZ (Taylor,

1992; Meffre and Crawford, 2001).

While it is largely agreed that these features are produced by the subduction of the

DEZ and WTM, there is disagreement regarding the morphology of the arc prior to collision. Some research provides pre-collisional profiles dominated by a high-standing forearc, suggesting that the arc was unusual even before collision (Daniel and Katz, 1981;

Green and Collot, 1994; Pelletier et al., 1994). Others suggest that nearly all the anomalous features of the arc are due to the collision, and that prior to collision the arc exhibited more conventional features (Taylor, 1992). Meffre and Crawford (2001) compared the New Hebrides arc to other western Pacific island arc systems undergoing collision, and found that they all shared a shallow trench and uplifted forearc, supporting the conclusions of Taylor (1992). These authors also suggested that arcs undergoing deformation due to collisions are able to revert to their original morphology relatively quickly afterwards, in about 1.5 – 3 Ma.

The tectonic history of Araki and central New Hebrides as a whole has been influenced by the subduction of these large features, which has caused variable uplift and subsidence. Corals from southern Espiritu Santo island provide evidence of subsidence prior to the current uplift of the forearc islands (Taylor et al., 2005). These corals, dated from 210 to 410 ka, along with their elevation and surrounding stratigraphy, led to the conclusion that central New Hebrides must have undergone hundreds of meters of subsidence followed by an abrupt shift to uplift. The shift from subsidence to uplift likely occurred over the time interval of 170 to 100 ka. The model used to explain this shift features a “crustal shortening” mechanism, where the impingement of a large object such

as a ridge or seamount causes horizontal shortening, leading to forearc thickening and uplift, as well as interplate coupling (Figure 2). When the impinging object breaks or becomes decoupled, the forearc extends and subsides. As more large features approach the plate boundary, the forearc begins uplifting again. There is evidence of previous similar vertical oscillations at this site throughout the Quaternary, indicating a cyclical process that occurs over a period of one to two hundred thousand years (Taylor et al.,

2005).

The more recent (~100 ka) tectonic history of the New Hebrides arc varies from island to island, though most islands are generally uplifting. Taylor (1992) observed that the presence of young, emerged coral at the top of a stratigraphic unit indicates tectonic uplift, whereas lack of terraces indicates stability or subsidence. Based on such patterns, it can be inferred that islands such as Malakula and Espiritu Santo (and by extension,

Araki, since it is very close to Espiritu Santo) have experienced various rates of uplift over the past 100 ka. The backarc islands, Pentecost and Maewo, have also experienced uplift (Taylor, 1992). In contrast, Aoba and Ambrym have no coral terraces and are subsiding about half as much as the forearc is uplifting (Meffre and Crawford, 2001).

1.2 Determination of past sea levels using coral terraces

Uplifted coral terraces such as those on Araki provide an opportunity to estimate past sea levels. The well-known coral terraces on Huon Peninsula (HP), which is on the northeast coast of Papua New Guinea, have been studied extensively (Chappell, 1974;

Ota et al., 1993; Chappell et al., 1996; Esat et al., 1999; Yokoyama et al., 2000). A discussion of the methods involved in determining past sea levels from Huon Peninsula terraces will provide insight into applying similar techniques to Araki. Additionally, HP and Araki are likely to have experienced similar glacio-isostatic histories (Chappell et al.,

1996) and therefore similar relative sea level histories. If they share a similar relative sea level history, then an HP-derived sea level history can be used as a benchmark when determining Araki uplift history.

Like Vanuatu, the Huon Peninsula is located on the overriding plate of a subduction zone. The tectonic environment of Papua New Guinea is complex, including at least two to five microplates trapped between the Australian and Pacific plates and extensive seismic activity (Ghasemi et al., 2016). The Huon Peninsula and the northern part of Papua New Guinea sit on the overriding South Bismarck plate, while the remainder of the country is on the Australian plate (Tregoning et al., 1998) (see Figure 3).

Multiple uplifted coral terraces on HP indicate a history of Quaternary uplift (Chappell,

1974). Since the Holocene there is evidence for coseismic uplift as well, though over longer timescales this becomes less significant (Ota et al., 1993).

Figure 3: Overview of Papua New Guinea and surrounding tectonic plates. Purple cross-hatching indicates terrain with an elevation above 1,000 m (Adapted from Robbins et al., 2013)

In addition to signaling past tectonic uplift, coral terraces can also be used to determine past sea levels. Chappell et al. (1996) deduced past sea levels for various transects along the coast, using several common assumptions in coral-based sea-level studies. One of the most important aspects of determining past sea levels is the designation of a reference terrace that can be used to determine a long-term uplift rate, which is then factored into calculations. For Huon Peninsula calculations, the terrace associated with the peak of the Last Interglacial (LIg) is used (Chappell et al., 1996). The

LIg occurred from about 118.5 to 129.5 ka (Hibbert et al., 2016) and had sea levels 6.6 ±

2 m higher than modern levels (Kopp et al., 2013; Dutton et al., 2015). The timing and

sea level for the LIg is well known and thus is a good choice for reference terraces. This is due to the higher sea levels, which allowed Last Interglacial terraces to be well recorded on stable coastlines, such as the Seychelles (Dutton et al., 2015), Western

Australia (Stirling et al., 1998), and the Bahamas (Chen et al., 1991; Thompson et al.,

2011). The Last Interglacial terrace on HP was recorded at elevations of 320 m or 403 m, depending on the transect. Chappell et al. (1996) chose an age of 122 ± 4 ka for the peak of the Last Interglacial, based on dating of the Last Interglacial high sea levels from a range of sites around the world. Further, they adopted a value of 5 ± 2 m above modern sea level as the estimated sea level at the time. Using these values, Chappell et al. (1996) were able to calculate uplift rates (U) for several transects, ranging from 2.8 to 3.3 mm/yr, using the following equation:

U = (H* – S*)/t* Eq. 1 where H* is the height of the reference terrace ( 320 m or 403 m), S* is the sea level at the time of reference terrace formation (5 m), and t* is the age of the reference terrace

(122 ka). They assumed that for any given transect, this uplift rate was constant over time scales exceeding several thousand years (Chappell et al., 1996). Corals from various points along each transect were dated, and the sea level (S) from the time of coral formation (t) could be calculated using:

S = (H + z) – Ut Eq. 2 where H is the present elevation of the coral, z is the growth depth, U is the tectonic uplift

(Eq. 1), and t is the age of the coral fossil.

At higher uplift rates there is more vertical difference between sea level events.

Further, corals deposited during times of low sea level are more likely to be found above

modern sea level and be better preserved if high uplift rates are present. The high uplift rate on the Huon Peninsula allows for a more complete sea level history of the last glacial-interglacial cycle to be established than is possible in locations with slower uplift rates such as Barbados in the West Indies (about 0.3 mm/yr) (Gallup et al, 2002; Cutler et al., 2003).

Chappell et al. (1996) assumed sea level on the Huon Peninsula would be minimally affected by hydro-isostatic and global gravitational effects because it is distal from ice sheets and lacks a continental shelf. However, Lambeck et al. (2002) determined that isostatic effects could not be ignored and used coral ages from Chappell et al. (1996), an assumed growth depth of 5 ± 5 m, and isostatic modeling to produce the sea level curve seen in Figure 4. On a global scale, Araki and Huon Peninsula are relatively close to one another. This proximity, combined with their distance from the continental ice sheets and similar levels of ocean loading, means Araki and Huon Peninsula are likely to have comparable hydro-isostatic histories and thus share a common sea level history

(Lambeck and Chappell, 2001). This shared history allows for a comparison between the estimated Vanuatu uplift history and the Huon Peninsula sea level record.

Figure 4: Ice-volume equivalent sea level (sea level associated with the change in ocean volume due to growth or melting of land-based ice sheets) curve derived from Huon Peninsula uplifted corals and sediments from Bonaparte Gulf (Lambeck et al., 2002).

1.3 History of fossil coral studies on Araki

As methods of surveying and dating corals have improved, so too has the knowledge gleaned from studies of Araki corals improved. Jozsef Urmos, a Masters student at Cornell University, conducted the earliest research on the island in 1985, detailing the various terraces and collecting corals (Urmos, 1985). Terrace heights were measured with an error of ± 2 m. Uranium and thorium quantities in corals were determined using alpha spectrometry, which has much larger error than modern mass spectrometry methods (typical errors for 35 – 105 ka ages are about ± 4 – 14 ka 2σ using alpha spectrometry, while mass spectrometry methods for the same age range have errors of about ±0.03 – 0.3 ka 2σ) (Gallup, 2015; Cheng et al., 2013). Urmos used the 230Th

dates of the terraces to identify periods of coral growth and estimate uplift rates. He concluded that the long-term mean uplift rate from 105 ka to present was relatively constant at 2.32 mm/yr. He also determined that the uplift rate has been nearly twice that since the Holocene, though the exact timing of that shift was not clear. While this research forms a good foundation, the development and improvement in mass spectrometric methods of U and Th isotopic measurements and higher precision GPS now allows for improved quantitative constraints on tectonic rates.

Following Urmos’ work, Araki was not the focus of scientific study for many years. Research instead focused on the New Hebrides arc as a whole and understanding the collisional tectonics of the area (Maillet, 1989; Taylor, 1992; Pelletier et al., 1998;

Meffre and Crawford, 2001). Drilling was conducted at areas near Araki (see section 4.3 below) (Cabioch, 2003; Cabioch et al. 2003) and the Urmos data were used for comparison. In 2004, a set of Urmos corals were re-dated using mass-spectrometric methods, although the purpose was to extend the radiocarbon calibration curve, and not to focus on the tectonic history of Araki (Fairbanks et al., 2005; Chiu et al., 2005). The samples used in the present study were collected in 2005 by Fred Taylor and Nick

Freiburger. Freiburger (2006) dated 12 samples, but did not interpret the tectonic significance of the data. Carlson (2008, unpublished data) dated 13 additional samples from the collection and focused on calcite content in corals as related to their age. The remaining 36 samples are dated and analyzed in this study.

1.4 230Th Dating

230Th dating is based on the decay of 238U into two of its daughters, 234U and 230Th

(see Figure 5 for the full 238U decay chain). The half-lives of 238U, 234U, and 230Th are 13

4.4683 ± 0.0096 billion years (Villa et al., 2016), 245,620 ± 260 years, and 75,584 ± 110 years, respectively (Cheng et al., 2013). Because the average half-life of 238U is orders of magnitude greater than those of 234U and 230Th, we can assume that an infintesimal amount of 238U is lost due to radioactive decay with respect to the total amount of 238U

(Dickin, 2018). This simplifies the calculation as we are able to assume that the current

238U concentration is approximately the same as it was when the coral first formed, and we can focus soley on the formation of 230Th as a chronometer.

Figure 5: Uranium series decay chain. Red box indicates isotopes used in 230Th dating. See text for more accurate half-lives of isotopes pertinent to this study. (Wikipedia Commons, 2014) 14

Uranium is highly soluble in water, while thorium almost immediately adsorbs onto particles, giving it a very low solubility (Dickin, 2018). When weathering of continental rocks occurs, uranium is much more likely to dissolve and travel through surface waters to the ocean, while thorium concentrations will be very low in most water bodies. As corals grow, they incorporate uranium into their skeletons at a fairly constant rate so that they contain 2 - 3.5 ppm (Gvirtzman et al., 1973; Shen and Dunbar, 1995), but have essentially no thorium. Thus, any thorium later measured theoretically came purely from the decay of 238U, and the ingrowth of 230Th provides a chronometer.

Nuclides in a closed-system environment left for a long period of time relative to their half lives enter into a state of secular equilibrium (Dickin, 2018). In secular equilibrium the rate of decay (or ‘activity’) of the daughter nuclide is equal to that of the parent nuclide, such that:

퐴푐푡푖푣푖푡푦 = 휆0푛0 = 휆1푛1 = 휆2푛2 = 휆푁푛푁 Eq. 3 where λ is the decay constant, n is the number of atoms, and the subscript indicates parent, first daughter, second daughter, etc. Thus, in secular equilibrium the activity ratio of any daughter to parent is one (e.g. λ234n234/λ238n238 = 1). This ratio can be disturbed during natural processes. For example in ocean waters the ratio of 230Th/238U is essentially zero, because U is highly soluble in water while Th is not. The decay of 238U to 234U occurs via alpha decay, wherein an alpha particle is ejected from the atom. This causes alpha recoil, which damages and weakens the chemical bonds surrounding the daughter atom (Cherdytsev et al., 1961). Due to this damage, 234U is weathered out of materials preferentially, giving the modern ocean an 234U/238U activity ratio of 1.147, corresponding to an enrichment of 14.7% relative to secular equilibrium (Stirling and

Andersen, 2009). This value is often presented in delta units, which expresses 234U as a a per mil value in reference to 238U, i.e.:

234 234 푈휆234 훿 푈 = [238 − 1] × 1000 = 147‰ Eq. 4 푈휆238

As decay continues, 230Th grows in while excess 234U decays. The age of a sample can be determined as a function of decay constants, the measured δ234U value, and the ratio of

230Th to 238U, as seen below:

230푇ℎ 훿234푈(0) 휆 − 휆 [ ] = −푒−휆230푇 + [ ] [ 230 234 ] + 1 Eq. 5 238푈 1,000 1− 푒−(휆230−휆234)푇 푎푐푡

and

훿234푈(푇) = [훿234푈(0)][푒휆234푇] Eq. 6

In these equations, δ234U(0) refers to the initial δ234U value and T refers to time (age). T is found by solving the equations iteratively (Edwards, 1988).

During glacial periods there is evidence that the marine δ234U value may fall as low as 135‰, as excess 234U may accummulate in frozen sediments or become trapped in isolated subglacial lakes, which then releases back into the ocean during deglaciations causing the marine δ234U value to rise back to 147‰ (Chen et al., 2016). There is no evidence to suggest that the 234U/238U activity ratio of seawater has been significantly higher than 147‰ over time (Henderson et al., 1993), meaning that calculating the initial

δ234U value can provide a quality check, as increased ratios in corals must have arisen from diagenesis.

1.5 Coral Sample Selection

Before beginning the process of chemically dating a coral sample, the sample can be visually assessed to determine its likelihood of providing accurate age results. First, the coral should be minimally altered. This is shown by closely resembling modern coral and being white in color (Edwards, 1988). If the sample looks lacey or fragile, it has likely been partially dissolved, which may involve mobilization of U-series isotopes that would result in an inaccurate age. Ages can also be shifted by cementation; if the sample has additional aragonite or calcite, the age may be incorrect, depending on when the cementation occurred (Scholz et al., 2004). Finally, it is useful to hold the sample in the light and look for glints of crystal faces. This is a good initial check for the presence of calcite and recrystallization (Freiburger, 2006). A more precise quantity of calcite can be determined in the lab using X-ray diffraction (XRD); the generally accepted goal is to have less than 2% calcite (Bloom et al., 1974), although calcite amounts below the detection limits of the machine are ideal. A coral meeting these specifications may still have been altered, but shows potential to be a good candidate for U-Th dating.

2. Methods 2.1 Surveying and Collection of Samples

Fred Taylor (UT Austin) and Nick Freiburger (UMD) collected coral fossils from the island of Araki, Vanuatu in 2005. They were collected by hand using a hammer and chisel from their growth positions. Elevations were determined using a TopCon surveying level loaned by the Vanuatu Surveying Department. They leveled from the

highest living corals and closed loops to other highest living corals, a method which involves creating a continuous closed circuit of sight lines in order to estimate and account for error. Loops were closed to within a few centimeters, a variation likely due mainly to the slight differences between one shallow coral to another. Because this method used a cascade of loops there is a potential for errors to be cumulative, so elevation error was conservatively estimated to be ± 1 m (Fred Taylor, personal communication). The majority of the coral samples collected were determined to be of the genus Porites, a stony coral common in the Indo-Pacific region. Porites corals are relatively versatile, as they are tolerant of a range of sea surface temperatures and can grow at a range of depths, commonly from 1 to 25 m below sea level (Pichon, 2011). A comprehensive study by Hibbert et al. (2016) used observations of over 50,000 Porites species corals and found that the majority grow within the first few meters, but can grow nearly 100 m deep. However, the median growth depth is 4 m and 95% of all observations were found at 0 – 66 m below sea level. Hibbert et al. (2016) further broke down Porites samples by location, and found that Porites growing in the Pacific Ocean tend to grow much shallower than those in the Atlantic, having a median growth depth of essentially zero (compared to an Atlantic median growth depth of 7.3 m below sea level).

2.2 Process of U-Th Dating

The process of isotopically dating coral samples can be divided into three major parts: physical processing, chemical separation, and mass spectrometry. Physical processing begins by breaking apart larger coral samples with a hydraulic press in order to access less altered areas of the sample. Samples are further crushed using a mortar and

pestle, which allows for the selection of the most pristine pieces using tweezers and a binocular microscope. The sample is placed in a labeled plastic vial, with an additional sample set aside for XRD analysis. All surfaces of the work area are wiped down in between samples to avoid contamination.

The chemical processing of the sample, as described by Edwards, 1988 and updated in Cheng et al., 2013, has the overall goal of separating the U and Th components in preparation for mass spectrometry. Because the quantities of U and Th are so small (ppm or ppb), the chemical preparation must be completed in a clean lab.

One of the most important steps in the chemical analysis is the addition of a known spike, in this case 233U – 236U and 229Th, to the unknown isotopes, 238U, 234U, and 230Th. The spike isotopes are used because they are not naturally occurring, and have no possibility of being in the sample. Addition of a spike makes two important calculations possible:

(1) isotope dilution, which allows for the determination of the amount of unknown based on the ratio of unknown to known, and (2) fractionation correction, where the determination of the fractionation during mass spectrometry of the known spike isotopes makes it possible to correct for fractionation in the unknowns. After this, an iron solution is added and the resulting solution is titrated with ammonium hydroxide (NH4OH).

During titration, the iron will cause actinides such as U and Th to precipitate out with

FeOH, while leaving the major cations (Na+, Ca2+, Mg2+, K+, Sr+) in solution. The solid precipitate is then separated from the liquid with a centrifuge, redissolved, and run through a stepwise anion exchange resin, which isolates and separates the U and Th components. Upon dissolving and diluting the U and Th separates, they are ready for mass spectrometry.

Analyses were conducted using a Neptune Plus Inductively Coupled Plasma

Multi Collector - Mass Spectrometer (ICPMC-MS), an instrument that allows for quantification of ions at extremely low concentrations (Thomas, 2013). The liquid sample is pumped into a nebulizer, which uses argon gas to convert the sample into an aerosol.

The aerosol travels through a spray chamber, which filters out the larger droplets. The remaining fine aerosol exits the spray chamber and is transported to a plasma torch, which converts the elements in the sample into positive ions. The ions are then guided into the mass spectrometer, which uses a magnetic field to deflect ions based on their mass. These ions are detected and measured using either Faraday cups for uranium or a secondary electron multiplier for thorium. Faraday cups are conductive cups which, when hit with ions, produce a small electric current which can be measured. This produces high-precision results, because the current is proportional to the strength of the ion beam.

Uranium isotopes are measured using this method, but thorium isotopes have much lower concentrations, and Faraday cups are not sensitive enough to measure them. Instead, thorium is measured using an electron multiplier, which is less precise due to the smaller number of ions measured, but more sensitive. In this method, an ion strikes a dynode, which causes the release of secondary electrons. These electrons hit another dynode, which then produce even more secondary electrons. In this manner, a single ion produces a shower of electrons collected by an anode (Edwards et al., 1986; Edwards, 1988).

For each sample, the following isotopes are measured: 233U, 234U, 235U, 236U,

229Th, 230Th, and 232Th. 238U is not directly measured because its ion beam is very large compared to the other U-series isotopes, but is instead determined from its ratio with

235U. The ratio of 238U to 235U is currently 137.88 on Earth and only changes very slowly because each of these isotopes is the parent of a long decay chain ending in a stable

isotope of lead (Dickin, 2018). In addition, one blank sample is run for about every ten samples and a standard is run every 3-4 samples. Raw data from blanks, standards, the spike, and the sample are imported into Excel. Isotope values and ratios are calculated, corrections for blanks, electronic background, the tail of the very large 238U beam when appropriate, and mass fractionation are made, and the 230Th age equation (see Eq. 5 and

6) is solved iteratively to provide sample ages, as well as 238U (ppm), 232Th (ppt),

230Th/232Th, δ234U (measured), δ234U (initial), and 230Th/238U values (see Appendix A).

Samples are evaluated for quality of 230Th ages based on the following criteria

(Broecker and Thurber, 1965): (1) lack of recrystallization of aragonite to calcite, (2) confirmation that bulk U concentrations are within range of modern samples of the same species, (3) a low concentration of 232Th and a high value of 230Th/232Th, indicating the presence of 230Th is due to in situ decay of the parent, and (4) 234U/238U values that yield back-calculated initial 234U/238U values within range of modern seawater. This study will accept a broader range of 234U/238U than just modern seawater because the sample set includes glacial periods, which have lower values of δ234U (see Results section).

2.3 X-Ray Diffraction

Powder X-ray diffraction (XRD) was conducted using a Philips X’pert Multipurpose

Diffractometer in the UMD Research Instrumentation Lab. To prepare the sample, fragments of coral are ground to a fine powder using a mortar and pestle. This powder is then sifted through a sieve into a sample holder, ensuring a randomized orientation of crystals. The sample holder is then inserted into the XRD machine, which directs X-rays onto the sample at a range of angles, from 5 ° – 65 ° (training by Tsutomu Shimotori,

UMD). As this happens, the angle and intensity of the reflected X-rays are recorded, and a diffractogram is produced (Figure 6).

Figure 6: Powder diffractograms. (a) and (b) are for calcite and aragonite standards. Note the strong peaks at 29.5° and 26° respectively (RRUFF Project). (c) Diffractogram for sample L in current study. Note how it is similar in appearance to the aragonite diffractogram, but includes a small peak at 29.5°, indicating the presence of calcite.

The diffractogram is typically analyzed using MDI Jade software and a Reitveld

Refinement, which calculates a weight percent aragonite in the sample. However, when calcite weight-percent standards were created and tested to verify sample results, it was discovered that this refinement was not working properly. It appeared as though the calcite in the standards was aligning itself into a preferred orientation, causing one peak in the diffractogram to read more strongly than others, and disrupting the curve-fitting software. As an alternative, a calibration curve was created using a series of standards with the following weight percent calcite: 0.1, 0.25, 0.3, 0.4, 0.5, 0.75, 1, 1.5, 2, 2.5, and

3% (Chiu et al., 2005; Sepulcre et al., 2009; Smodej et al., 2015). Duplicates were run for the 0.5, 0.75, 1, and 2% samples, which were averaged. The amount of calcite and aragonite in the powder mixture is generally proportional to the relative intensity of each component (Chiu et al., 2005). Relative intensities were calculated using peak areas from a representative peak of aragonite and calcite. The peak-area ratio (AR) for each standard was calculated according to the following equation:

퐴푅 = 퐴퐶/(퐴퐶 + 퐴퐴) Eq. 7

where AC is the area of the calcite peak at 2θ = 29.5° [104] and AA is the area of the aragonite at 2θ = 26° [111] (Milliman, 1974; Chiu et al., 2005; Smodej et al., 2015). No calcite was detected for the 0.1% standard, which is consistent with the approximate detection limits of ~0.2% in recent XRD calibration studies (Chiu et al., 2005; Smodej et al., 2015). Peak-area ratios were plotted against the weight percent and were fitted with a linear trendline, which is suitable for all calcite contents below 5% (Sepulcre et al., 2009;

Smodej et al., 2015). The trendline can be described by the equation y = 0.14±0.08x +

0.02±0.12, which has a p-value of 0.002 (see Figure 7). This line can be used to determine the calcite contents of the remaining samples, which range from 0 to about

2.5% calcite.

Figure 7: Calibration curve used for determination of aragonite content. A linear trendline is used to describe the relationship between weight percent calcite and the peak area ratio.

3. Results

Data for thirty-six samples are presented in this study, which range in age from

32.16 ± 0.04 to 141.3 ± 0.3 ka. They were collected from Araki at elevations ranging from 39 ± 1 to 180 ±1 meters above sea level. Calcite content ranges from about 0 to

2.5% ± 2% (Appendix B). In addition to calculated age and calcite content, the following values are also presented for each sample: 238U and 232Th concentrations, δ234U values, and 230Th / 232Th and 230Th / 238U isotopic ratios (see Appendix A). According to

Yokoyama et al. (2001), an acceptable range of 238U is 2.5 – 3.5 ppm based on average U

concentrations of corals in the Pacific, and is weakly dependent on the species of coral.

Samples in this study range from 1.3 – 3.7 ppm, though there are a significant number near 2 ppm. Given that the samples with 2 ppm U are not anomalous in any other way, only Sample E, which has a U concentration of 1.3 and the lowest 238U value, is excluded. Pristine corals from oceanic islands tend to have less than 0.5 ppb of 232Th

(Edwards et al., 1986); all samples presented here have 232Th values significantly less than this.

As previously mentioned, δ234U values can also provide a quality check, as samples are expected to have similar values to that of modern seawater (147 ‰) or 5-10

‰ lower than modern seawater during glacial periods, such as 50-30 ka (Esat and

Yokoyama, 2006). Values higher than the modern seawater value tend to correlate with older apparent ages (Gallup et al., 1994). δ234U values for our samples range from 134‰ to 153‰. For our purposes we will use a range with a more conservative upper limit, as higher values typically correlate to diagenesis. A larger range for the lower limit is used because our samples include glacial periods. Therefore, the acceptable range of δ234U values is within error of the range of 147‰ + 8‰ and – 10‰; samples outside this range are assumed to have experienced diagenesis. Calcite content of less than 2% is typically acceptable (Bloom et al., 1974; Edwards, 1988; Yokoyama et al., 2000; Cabioch, 2003;

Dutton et al., 2015); however, because the error associated with the calibration curve is

~2% for most samples here, all those with a measured calcite value greater than zero are excluded. Following these exclusions, the overall age and elevation range remains unchanged.

The samples dated here expand upon the dataset of Araki corals begun by Nick

Freiburger in 2006 and Tyler Carlson in 2008. They previously dated twenty-three samples, which, when added to the current set comprise a dataset of fifty-nine samples.

Of these, there are forty-one samples which meet the criterion for acceptable corals

(δ234U value of 147‰ + 8‰ and – 10‰ and calcite content of 0%), maintaining the age and elevation ranges of 32.16 ± 0.06 to 141.27 ± 0.3 ka and 32 ± 1 to 183 ± 1 meters above sea level, respectively (see Appendices A and B for full datasets).

Although there is a wide range of growth depths of Porites corals (Pichon, 2011),

Hibbert et al. (2016) found a median value of 4 m. Lambeck et al. (2002) assumed a growth depth of 5 ± 5 m for Papua New Guinea corals (see section 4.5 for more details and error analysis). As the Lambeck et al. (2002) sea level record is optimal for comparing with the Araki data, it is appropriate to adopt the same assumed growth depth of 5 ± 5 m for all Araki corals.

4. Discussion

With the addition of these new coral specimens, it is possible to better resolve the tectonic history of Araki. As described in sections 1.1 and 1.3, previous work documents:

1) early subsidence (410 to 210 ka) that shifted to uplift by 170 to 100 ka; 2) a long-term average uplift rate from ~105 ka to present, established by Urmos (1985); and 3) a rapid uplift rate during the Holocene, which is approximately double the long-term rate and appears to extend back to the Last Glacial Maximum (Taylor et al., 2005; Cabioch et al.,

2003). This discussion will explore early subsidence on Araki, how far back the

Holocene rate may extend, and attempt to determine a feasible uplift rate prior to the 27

Holocene. The predicted uplift rates will be used to infer a sea level history, utilizing

Equation 2 as presented in Section 1.2. This sea level history can then be compared to the Huon Peninsula sea level record; the degree to which the sea level records agree gives confidence to the model.

4.1 Early subsidence from 125 to 106 ka

As previously mentioned, New Hebrides experienced hundreds of meters of subsidence prior to the current uplift. Data presented here from the Last Interglacial (125 ka) allows for the shift from subsidence to uplift to be constrained for Araki.

On a stable coastline, it would be expected that a reef formed during the Last

Interglacial (LIg) would be ~21 m above a reef formed at 105 ka. This is because sea level during the LIg was 6 m higher than modern levels (Cutler et al. 2003), and sea level at 105 ka was 15 m lower than modern levels (15 m + 6 m = 21 m difference) (Bloom et al., 1974). However, this is not what was found on Araki.

Sample BP was formed at the height of the Last Interglacial and is dated as 125.2

± 0.8 ka. It has an elevation of 179 m. Urmos (1985) collected two coral samples from the summit of Araki at an elevation of 233 m and dated them as 104 ± 11 ka and 107 ± 8 ka (2σ). For the purposes of this study, we will take a weighted average of the samples and use an age of 106 ± 6.5 ka for the uppermost terrace of Araki. This means that as opposed to the LIg reef being ~21 m above the 105 ka reef, it is 54 m below, and 75 m lower than we expect (54 m plus the 21 m due to sea level change). Therefore, subsidence continued from 125 ka to about 106 ka at about 3.9 mm/yr (75 m / 19 ka), after which uplift began, perhaps abruptly.

4.2 Long-term uplift rate

The average uplift rate for the last ~106 ka is based on samples from the summit of Araki. As previously mentioned, Urmos (1985) collected two coral samples from the summit of Araki at an elevation of 233 m, whose age is a weighted average of 106 ± 6.5 ka. Papua New Guinea data from Bloom et al. (1974) indicates that sea level was about

15 m lower than modern at 106 ka, meaning Araki must have uplifted 15 m in addition to the 233 m known from the top terrace. Another way of explaining this data point is by visualizing the uplift of the equivalent of the modern coastline over the past 106 ka. As corals grew when sea level was 15 m below the elevation of modern sea level and they are now at 233 m, this means the equivalent of the modern sea level 106,000 years ago has uplifted a total of 248 m in the interim (this time and elevation will be used in the model below). The uplift rate can then be calculated as: 248 m/106 ka = 2.34 mm/yr

(green line in Figure 10 and 11).

4.3 Uplift rate since ~20 ka

In Araki, Holocene terraces are higher than expected based on the heights and ages of older terraces. The top of the

Figure 8: Locations of drilling locations with reference to Araki. Scale Holocene reef on Araki is 28 m bar denotes 2 km. Location of Araki as seen in Figure 1. (Google and ZeeMaps) above living corals and is about

7 ka old (Taylor et al. 2005). Given that eustatic sea level was 3-5 m below present at 7

ka (Fleming et al. 1998), if the long-term average uplift rate of 2.34 mm/yr was in effect since the Holocene, the reef would have risen about ~16 m (2.34 mm/yr x 7,000 years) and would be at ~12 m above modern sea level today. This implies that the rate of uplift accelerated at some point prior to the Holocene. To determine the uplift since 7 ka, the amount of uplift is divided by the age of the coral. As Araki uplifted 31 – 33 m (28 plus the 3-5 m it was below modern sea level initially) and it is 7 thousand years old, this corresponds to a rate of ~ 4.6 mm/yr for the Holocene (orange line in Figure 10 and 11).

Drilling at two locations near Araki adds additional constraints on the uplift rate back through the Last Glacial Maximum (~20 ka). Drilling was conducted at

Tasmaloum, a coastal area on Espiritu Santo to the west of Araki (Cabioch, 2003), and at

Urelapa, an island to the east of Araki (Cabioch, et al. 2003) (see Figure 8). Holocene- dated (6 ka) corals were found at 35 m above sea level on Tasmaloum, and it was estimated that 2 m of that uplift was due to hydro-isostatic effects (Taylor et al., 2005).

The corals suggest an uplift rate of about 5.5 mm/yr for Tasmaloum (Taylor et al., 2005).

At Urelapa, samples aged about 6 ka and 23 ka were found about 51 m apart, corresponding to an uplift rate of around 3 mm/yr (Cabioch et al. 2003). Because LGM- aged corals were found at Tasmaloum and Urelapa at 19-28.5 m and 50-55 m below modern sea levels, respectively, it is fair to assume that uplift rates have been at least this high for the last ~20 ka. At 20 ka sea level was at least 120 m below present sea level

(Lambeck et al. 2002), meaning that the corals have uplifted significantly since then

(from -120 m up to the elevation at which the corals were found). The corals’ uplift over

20 ka corresponds to rates of 4.6 – 5.0 mm/yr for Tasmaloum and 3.25 to 3.5 mm/yr for

Urelapa. Urelapa is farther away from the trench, and is therefore likely to experience

less uplift due to the subduction of the WTM and DEZ than the more western islands. As

Araki and Tasmaloum are closer to the trench, their higher uplift rate estimations of 4.6 –

5.0 mm/yr are consistent with their tectonic setting. Additionally, Araki’s position between Tasmaloum and Urelapa suggest that Araki’s Holocene uplift rate extends back through at least 20 ka as well.

4.4 Modeling uplift rate from 20-105 ka

To determine how far back in time the higher Holocene uplift rate extends, it is necessary to create a model of Araki’s uplift history. This model can be visualized on a plot of elevation versus time (Figure 9).

Figure 9: Simplified model of Araki uplift history. Lines represent uplift incurred over time, and their slope equals the uplift rate. Hol represents the faster Holocene uplift rate, LT represents the slower long-term average uplift rate, and Pre-LGM refers to the hypothetical rate that must occur prior to the Last Glacial Maximum in order to accommodate both Hol and LT rates. A is the point at which the island switches from subsidence to uplift, at 106 ka, a point that remains constant in this model. B is the point at which the Pre-LGM rate shifts to the Hol uplift rate, a point that is adjusted in this model. Previous work provides a long-term uplift rate of 2.34 mm/yr over the last 106 ka

(Urmos, 1985), represented by the line labeled LT in Figure 9. Drilling at Tasmaloum

(Taylor et al., 2005) and Urelapa (Cabioch et al., 2003) suggests rapid uplift since at least

20 ka, shown as the steeper line Hol in Figure 9. To give an overall rate of 2.34 mm/yr while also including a faster uplift rate of 4.6 mm/yr for at least 20 ka, the pre-LGM rate must be less than, or a shallower line than, 2.34 mm/yr. Note that a single pre-LGM uplift rate is the simplest assumption for the tectonic history, although there is the possibility that it was not uniform. However, right now there is no firm justification for a more complex tectonic history, so the simplest assumption will be used. Determining the Pre-

LGM rate and the timing of the rate change involves first fixing the pre-LGM rate to assume it began at 106 ka (Point A in Figure 9 and 10). Then, by changing the assumption of how far back the Holocene uplift rate goes (Point B in Figure 9 and 10), various uplift histories are created. The success of an uplift history comes from combining it with the ages and elevations of Araki fossil coral samples to generate a sea level record, as discussed in Section 1.2 and modeled in Figure 10.

Constructing the model involves plotting the fixed point at which the island switches from subsidence to uplift (A), and the adjustable point at which the Pre-LGM rate switches to the Holocene rate (B). The slope of the line connecting these two tie- points is the Pre-LGM uplift rate. Point A, when described as an (x, y)-point, needs both an elevation and a time. It has been previously determined that the shift occurred at or before 106 ka. Using the long-term uplift rate of 2.34 mm/yr, the island would have uplifted 248 m, thus Point A is (106 ka, 248 m). Next is Point B, representing the shift from Pre-LGM rates to the Holocene rate, an assumption that can be adjusted in future iterations of the model. There is evidence that the Holocene uplift rate extends back at least through 20 ka, so that will be used first. If the Holocene rate of 4.6 mm/yr occurred

for 20 ka, the island would have uplifted 92 m such that a coastline that had a sea level the same as modern at 20 ka would be at 92 m today, giving Point B the coordinates (20 ka, 92 m). The line connecting these points has a slope (and therefore uplift rate) of 1.81 m/ka, or 1.81 mm/yr (purple line in Figure 10 and 11). The vertical distance between coral samples and the Pre-LGM line represents the sea level at the time of coral formation plus the growth depth of the coral, previously chosen as 5 ± 5 m. This value can also be calculated using Eq. 2.

Figure 10 provides a visual description of the model using Sample AF as an example. Sample AF is dated as 45.5 ka and was found at an elevation of 70 m. Given this information, combined with the uplift rates described in the previous paragraph, all the components of Eq. 2 are available.

S = (H + z) – Ut Eq. 2

The value of H is the elevation of the coral, or 70 m, and the growth depth (z) is assumed to be 5 ± 5 m for all samples. The Ut term represents the total uplift experienced by the island since the formation of the corals, which is 137 m. This is the sum of the uplift experienced due to the Holocene uplift rate (4.6 mm/yr * 20 ka = 92 m) and the uplift due to the Pre-LGM uplift rate of 1.81 mm/yr (1.81 mm/yr * 25 ka = 45 m). Given these values, the sea level (S) at the time of coral growth (45.5 ka) can be calculated as 62 m lower than present (Figure 10).

Figure 10: Graphical explanation of the uplift model as a plot of elevation versus time showing how Eq. 2 is used to determine paleo-sea level, given an estimated Pre-LGM uplift rate, and sample AF as an example.

The next step is comparing the sea levels determined from this model to other, nearby sea level curves, such as Lambeck et al. (2002). With tie-point B at 20 ka (Figure

11), Figure 12a shows this produces calculated sea level values from the Araki corals of about 65 to 85 m below modern levels at 30 – 40 ka. Lambeck et al. (2002) predicts sea level at that time to be between about 70 and 95 meters lower than modern levels, values close to the projection. The model also calculates sea levels to be 45 - 50 m below present at 50 ka, slightly shallower than Lambeck et al.’s range of 50 – 65 m. The calculated sea level for 60 ka diverts more drastically from the Huon Peninsula curve, predicting about

38 m below present, while HP predicts lower sea levels, 55 – 65 m (see Figure 12a).

The initial assumptions can be adjusted to see if it makes the sea level estimations any better, by altering the timing of Point B (when the shift to a faster uplift rate occurred). For example, the uplift change can be pushed earlier, such as to 25 ka. Here the two tie-points are B = (25 ka, 115 m) and A = (106 ka, 248 m), which produce a Pre-

LGM line with a slope of 1.64 (red line in Figure 11). This rate brings predictions from

55 – 65 ka closer to the HP values, but they are still much too shallow. The new predicted sea levels for 30 – 40 ka are deeper, about 75 – 95 m, closer to the range provided by HP

(Figure 12b). Pushing the shift in uplift rates to be even earlier, at 30 ka, the light blue

Pre-LGM line is produced with a slope of 1.45 mm/yr. This line creates the highest projected elevations of the current coastline, predicting much lower paleo-sea levels.

These lower sea levels fit the HP curve very well in the 55 – 65 ka range, but are much too deep for 30 – 40 ka (Figure 12c). Adjusting in smaller increments than ~5 ka produces results with such little variance that it is likely rendered insignificant by error

(see navy line in Figure 11, and Figure 12d).

Figure 12 shows the calculated sea levels for various uplift rates overlaid on the sea level curve produced by Lambeck et al. (2002). It is obvious that there is some scatter, but the predicted sea level generally follows and matches the Lambeck et al.

(2002) curve. The good correspondence between the Huon Peninsula and inferred Araki sea level records provides important support for the validity of the Huon Peninsula sea level record, as there is very little fossil coral data worldwide for sea level between 70 and 20 ka. It appears as though higher rates (1.64 to 1.81 mm/yr) produce sea level elevations that fit the curve better from 30 to 55 ka, while the lower rate of 1.45 mm/yr seems to correlate better for 55 to 70 ka. This could be indicative of instances of

coseismic uplift around 55-60 years or of a change in uplift rate, although there isn’t strong evidence to signify another shift in uplift. The highest uplift rate, predicted from a shift to faster uplift at 20 ka, produces sea level elevations that fit the Huon Peninsula sea level record very poorly, suggesting that the shift in tectonic rate occurred some time between 30 and 25 ka.

The model presented here provides good evidence that Araki experienced a rate of uplift of 1.45 – 1.64 mm/yr starting at 106 ka, followed by a shift to a higher rate some time between 30 and 25 ka. This is a change in rate of nearly 3 mm/yr in a relatively short

(~5 ka) amount of time. This is in addition to the dramatic tectonic shift from subsidence to uplift just prior to 106 ka, and gives more evidence as to the extensive effects possible from the subduction of large seamounts, such as the WTM and DEZ. It could indicate that the down-going plate has coupled more strongly with the overriding plate in the last

25 ka, causing increasing uplift and deformation of the forearc.

Figure 11: Various uplift rates, in mm/yr, as well as coral samples used in this study. Vertical and horizontal error bars for coral samples are not visible at this scale. Vertical distance between coral elevations and lines of modern coastline elevation trajectories represents the sum of the paleo-sea level at that time plus the growth depth of corals (assumed to be 5 ± 5 m).

a. b.

c. d.

Figure 12: Calculated paleo-sea levels using various parameters, overlain on sea level curve from Lambeck et al. (2002). Colors correspond to Figure 11. (a) Shift to higher uplift rates occurs at 20 ka, rate = 1.81 mm/yr (b) Shift to higher uplift rates occurs at 25 ka, rate = 1.64 mm/yr (c) Shift to higher uplift rates occurs at 30 ka, rate = 1.45 mm/yr (d) Shift to higher uplift rates occurs at 22 ka, rate = 1.74 mm/yr. 38

4.5 Errors encountered and future work

The largest error encountered with this method is from assuming that corals grow at a depth of 5 m, when in reality they can grow up to 25 m below sea level (IUCN Red

List, 2018). If the model presented here were run again with the assumption that all samples grew at a depth of 25 m, the predicted sea levels would be shallower. This would suggest a slower uplift rate, and an earlier shift to the faster uplift rate (see Figure 13).

Figure 13: Predicted sea levels using various rates, assuming corals grew at a depth of 25 m below sea level. Colors correspond to lines in Figure 11. Overlain on sea level curve from Lambeck et al. (2002). This error would be better quantified for each coral if a coral biologist were to identify their species so that they could be matched up with a probability distribution depth for that species (Hibbert et al., 2016). Alternatively, if co-sampled coralline algae were used to constrain the depth at which corals grew, the error might be greatly reduced, as some algae have very narrow ranges of growth depth (Cabioch, 1999). Knowing the

approximate growth depth would improve the accuracy of calculated sea levels and allow more confidence in comparison with the reference sea level curve.

Another major source of error is the value used for the age of the peak elevation of Araki (104 ± 11 ka and 107 ± 8 ka), which is used as a tie-point (A) for generating slopes. This data point was determined in 1985 and was found using alpha spectroscopy, which gives a large error. If this sample were to be verified using high-precision ICP-MS methods, error could be reduced. The scale of error incurred by this data point can be visualized using the maximum and minimum values of 106 ± 6.5 ka as alternate tie- points in order to determine error. This involves creating 2 additional lines, both keeping the lower tie point (B) of (25 ka, 115 m), but changing the higher tie-point (A) to reflect the error in Urmos’ data (i.e. (99.5 ka, 248 m) and (112.5 ka, 248 m)). The difference between the original line and these new lines provides the estimate in error. Figures 14 and 15 show the error associated with the uplift rate of 1.64 mm/yr (red line in Figure 11) due to error in the older tie-point. Because the younger tie-point (25 ka, 115 m) remained the same, the error increases along with age (Figure 14). As such, since the samples are relatively young (less than 70 ka), they have fairly small errors (~ ± 5 m) (Figure 15).

Figure 14: Slopes (uplift rates) determined by using the minimum and maximum of 106 ± 6.5 ka, showing the error due to Urmos’ data. Note the increasing error with age.

Figure 15: Error incurred from use of maximum and minimum values of 106 ± 6.5 ka when determining sea level using a rate of 1.64 mm/yr. Overlain on sea level curve from Lambeck et al. (2002). A future endeavor that would benefit this analysis would be the establishment of a range of uncertainty inherent in the model based on the parameters involved. This could be done using a Monte Carlo approach wherein values for model input variables are randomly chosen from distributions based on known or estimated uncertainties. The main sources of error to be included in this analysis would be the growth depth of corals, the timing of the shift from subsidence to uplift (currently based on Urmos (1985) data), analytical uncertainty of ages, the elevation range of the corals from the field, and tectonic non-linearity. According to Hibbert et al.

(2016), 95% of Porites found in the Pacific Ocean grow at depths of 0 to -71 m, but

with a median growth depth of virtually zero. Depth distribution histograms, such as

Figure 16, could help with the accuracy of this portion of the analysis.

The error associated with Urmos’ data point (1985) has previously been discussed; however, there is also some error in the assumption that 106 ka was the point at which subsidence switched to uplift. That assumption was made because 106 Figure 16: Depth distribution histogram from Porites ka is the age of the peak of Araki. There sp. in Pacific Ocean (Hibbert et al., 2016) simply isn’t any data to suggest an earlier switch, although the possibility cannot be ignored. The next variable to be factored into the model is the analytical uncertainty of the fossil ages. The errors provided are instrumental error, but analytical error, as well as the long-term precision of the machine must also be taken into account (Cheng et al., 2013). The conservatively estimated +/- 1 m error from field surveying should also be incorporated into the model, which should also address the fact that although efforts were made to measure elevation based on the highest currently living coral, there may have been a higher coral elsewhere. Finally, the model is based on the idea that the tectonic motion of the island is divided into linear segments of constant rates, when in reality these rates were likely varied over time. In addition, periods of coseismic uplift or subsidence have been assumed to be insignificant over long periods of time, which may be a faulty assumption. The input of all of these sources 43

of error into a Monte Carlo analysis would generate an uncertainty range for the estimate of the Pre-LGM uplift rate and tectonic history of Araki, which would improve the confidence in the results of this study.

5. Conclusions:

A collection of 59 coral samples from Araki, Vanuatu and their corresponding elevations and 230Th ages are presented. Modeling based on these samples and previous knowledge of Araki helps to constrain the somewhat unusual tectonic history of Araki.

Prior to the Last Interglacial, Araki was experiencing subduction. Based on the findings of this research, the island continued to subside from 125 ka to 106 ka at a rate of about

3.9 mm/yr, at which point tectonic motion converted to uplift. From 106 ka to about 25 –

30 ka, Araki was uplifting at a rate of 1.4 to 1.6 mm/yr. After 25 - 30 ka, there was a shift to a higher uplift rate of about 4.6 mm/yr, a rate that continues to the present.

Further constraining the tectonic history of Araki helps to clarify how intraoceanic island arcs react when experiencing collision with large underwater features, an area that can be difficult to study due to evidence being subducted or otherwise obscured. The increased uplift rate in the last 25 – 30 ka could indicate that the downgoing plate has coupled more strongly with the overriding plate in that time, causing increasing uplift and deformation of the forearc.

While the focus of this study was on the tectonic history of Araki, the sea level record produced is also significant. In particular, it is valuable because it provides more information about sea level from about 50 – 70 ka, a period that is largely limited to

Huon Peninsula data. The strong match between Araki and Huon Peninsula sea level records provides support for the veracity of the Huon Peninsula record.

The conclusions of this study include an element of error that is difficult to quantify and could have significant impacts on the results: the unknown growth depth of the coral samples. Efforts to determine the growth depth of the coral samples or to re- date the samples collected from the 106 ka reef (Urmos, 1985) would do a great deal to improve and expand on the knowledge gained from this research. Nonetheless, the dataset presented here and the conclusions gleaned from it are valuable, providing information about a relatively unrecorded period of sea level history and evidence for uncommonly rapid shifts in tectonic motion.

45 References:

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50 Appendix A: 230Th Dataset

*Values were <0, indicating the blank signal was higher than the sample signal a Sample dated by Nick Freiburger (2006), b sample dated by Tyler Carlson (2008), samples dated by both are averaged Elevation error ± 1 m Growth depth = 5 ± 5 m for all samples -1 Calculations are described in section 1.4, the following decay constants are used: λ230 = 9.1705E-06, λ234 = 2.8221E-06, λ238 = 1.5513E-10 yr Acceptable range of δ234U for this study is 147‰ + 8‰ and – 10‰; 238U values > 1,900 ppb are acceptable as well. Highlighted samples are outside of acceptable ranges and are excluded from final analysis.

Appendix B: XRD Results

Sample Calcite Wt %* Sample Calcite Wt %* Number (err) Number (err) G 0.0% 2.4% AR 0.0% 2.4% AA 0.0% 2.4% AS 0.0% 2.2% W 0.0% 2.4% AP 0.1% 2.0% Q 0.0% 2.4% AQ 0.0% 2.2% R 0.0% 2.4% AT 0.0% 2.2% K 0.0% 2.4% AW 0.0% 2.4% P 0.0% 2.2% AV 0.1% 2.0% S 0.0% 2.4% BA 0.3% 2.0% A 1.2% 4.4% AM 0.0% 2.4% L 0.2% 2.0% AZ 0.0% 2.4% M 2.6% 9.6% BB 0.0% 2.4% N 0.1% 2.0% BP 0.0% 2.4% C 0.5% 2.3% BH 0.0% 2.4% O 0.6% 2.4% BE 0.0% 2.1% E 0.6% 2.4% BQ 0.0% 2.4% B 0.2% 2.0% BR 0.0% 2.4% X 0.0% 2.4% BO 0.0% 2.4% J 0.6% 2.6% BL 0.0% 2.4% H 0.0% 2.0% BD 0.0% 2.4% Y 0.2% 2.0% BJ 0.0% 2.4% Z 0.0% 2.4% BK 0.0% 2.4% F 0.0% 2.4% BN 0.0% 2.4% AB 0.1% 2.0% BC 0.0% 2.2% AC 0.0% 2.4% BM 0.0% 2.2% T 0.1% 2.0% * Calcite content determined using calibration curve AE 0.0% 2.4% described in section 2.3. XRD conducted on a Philips X’pert Multipurpose AG 0.0% 2.4% Diffractometer AD 0.1% 2.0% Highlighted samples are excluded from final analysis AH 0.0% 2.4% AK 0.0% 2.2% AF 0.0% 2.2% AL 0.0% 2.4% AN 0.0% 2.4% U 0.3% 2.0% AO 0.0% 2.4% 54