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NEW CHEMISTRY AND FROM ALKALINE IN THE NORTHWEST ROSS SEA, : INSIGHTS ON GENESIS ACROSS RIFTED CONTINENTAL AND OCEANIC LITHOSPHERE

Susan R. Krans

A Thesis

Submitted to the Graduate College of Bowling Green State University in partial fulfillment of the requirements for the degree of

MASTER OF SCIENCE

August 2013

Committee:

Kurt Panter, Advisor

John Farver

Chad Deering

© 2013

Susan Krans

All Rights Reserved

iii

ABSTRACT

Kurt Panter, Advisor

The West Antarctic Rift system hosts one of the world’s most extensive alkaline igneous provinces. Rifting that to the breakup of the proto-Pacific margin of

Gondwana in the Late Cretaceous was initially amagmatic and not related to mantle plume activity. Mantle heterogeneities and magmatic processes along a continent-ocean transect in the Antarctic Northwest Ross Sea region are deduced from geochemical and isotopic study of alkaline basalts. Specifically, new mineral chemistry and oxygen isotopes from the least differentiated basalts expand upon previous studies that have focused primarily on whole- data.

Alkali and basanite represent two end-members in the

Northwest Ross Sea and are characterized by high whole rock Mg# (59 ± 9), Ni + Cr (>

200 ppm), highly variable CIPW normative (> 20% hypersthene to > 20% ) and trace element contents (e.g., Sr = 400 to 1100 ppm, La/Yb = 11 to 28, Nb/Y = 1.2 to 3.6).

Phenocryst phases are primarily olivine (Fo66 to Fo91) and clinopyroxene (diopside) with rare amphibole (kaersutite) and exhibit varying degrees of compositional zoning.

Temperature and pressure estimates based on mineral-liquid equilibrium range from

1206-1331°C (olivine), 1220-1284°C and 0.9-1.3 GPa (clinopyoxene), indicating that early crystallization occurred at or below the Moho. Olivine oxygen isotopes measured by SIMS range from 4.71 to 5.44‰ and average 5.15 ± 0.52‰ and encompass values for clinopyroxene (4.97 ± 0.36 ‰) and (5.17 ± 0.10 ‰) measured by laser fluorination.

iv Correlation between oxygen isotopes and degree of partial melting (i.e. Nb/Y) suggests that lower degree melts preferentially consume a lower δ18O source interpreted as metasomatic veins in the lithospheric mantle. Temperature and pressure estimates across the continent-ocean transect indicate a region of lithospheric necking previously identified for Northern . Evidence for disequilibrium observed texturally and compositionally in suggests complex crustal processes, mainly in the ocean- continent transition zone. The results of this study support previous suggestions that late

Cenozoic alkaline magmatism in the West Antarctic Rift System is controlled by variations in partial melting of a heterogeneous mantle source and highlights future potential to investigate physical controls on volcanism at magma-poor rift margins.

v

To my father,

who instilled in me a sense of adventure, a love for science, and an eagerness to

learn.

To my mother,

who taught me to question everything and never be satisfied with “good enough”.

To my mentors and peers,

whose endless support and guidance has been a guiding light in the dark allies of

the unknown. I would not be where I am today if not for your wisdom and

comradery.

vi

ACKNOWLEDGMENTS

I would like to thank Noriko Kita and Jim Kern at the University of Wisconsin-

Madison SIMS laboratory, and Gordon Moore at the University of Michigan EMAL laboratory for technical support. I would also like to thank my committee members Dr.

Chad Deering (University of Wisconsin-Oshkosh, Dr. John Farver and especially my advisor Dr. Kurt Panter for all their guidance and support throughout this project. This research was funded by a collaborative grant NSF ANT-0943274.

vii

TABLE OF CONTENTS

Page

1. INTRODUCTION ...... 1

2. GEOLOGIC BACKGROUND ...... 5

3. METHODS ...... 7

3.1. Sample Summary ...... 7

3.2. Sample Preparation ...... 9

3.3. Mineral Chemistry ...... 10

3.4. Oxygen Isotopes...... 11

4. RESULTS ...... 12

4.1. Petrography ...... 12

4.2. Mineral Chemistry ...... 14

4.3. Oxygen Isotopes ...... 17

4.4. Thermobarometry and Estimates ...... 20

5. DISCUSSION ...... 25

5.1. Magma Origin ...... 25

5.2. Variations Across the Ocean-Continent Transect ...... 30

5.3. Crustal Influences on ...... 33

6. CONCLUDING STATEMENTS ...... 37

REFERENCES ...... 39 viii

LIST OF FIGURES

Figure Page

1 Map of NWRS samples ...... 49

2 TAS Diagram ...... 50

3 Glomercrysts and in thin section ...... 51

4 Disequilibrium textures in thin section ...... 52

5 Quench textures in thin section ...... 53

6 Mineral chemistry by zone ...... 53

7 Olivine chemical transects ...... 54

8 δ18O olivine vs. Fo ...... 56

9 Δ18O fractionation for NWRS mineral pairs ...... 57

10 Core-rim variation in calculated KD(Fe-Mg) exchange equilibrium ...... 58

11 Mineral-liquid tests for equilibrium ...... 59

12 Al-, Ti-Tschermak substitution in amphibole ...... 60

13 Amphibole P-T-H2O and thermal stability ...... 61

14 Differentiation trends ...... 62

15 Ni/Mn ratios for high Fo olivine ...... 63

16 Partial melting trends ...... 64

17 Degree of partial melting (Nb/Y) vs. δ18O olivine...... 65

18 T-P-Depth estimates for NWRS minerals ...... 66

19 T-olivine with distance from the CRS ...... 67

20 Clinopyroxene T-P with distance from the CRS ...... 68

21 Geochemical variations with distance from the CRS ...... 69 ix

22 Assessment of contamination from continental lithosphere ...... 70

23 Conceptual model for NWRS magmatism from continent to ocean ...... 71

x

LIST OF TABLES

Table Page

1 Whole-rock major and trace elements, and radiogenic compositions ...... 73

2 Petrographic description of samples ...... 76

3 Representative olivine chemistry ...... 78

4 Representative clinopyroxene chemistry ...... 83

5 Representative amphibole chemistry ...... 86

6 Representative chemistry ...... 87

7 Oxygen isotope composition of minerals ...... 88

8 T-P-H2O estimates ...... 89

9 Comparison of least differentiated samples by zone ...... 91

1

1. INTRODUCTION

The West Antarctic Rift System (WARS) hosts one of the most extensive alkaline igneous

provinces in the world, rivaled only by the East African Rift System. But unlike the East African

Rift, where rifting and magmatism has been attributed to active spreading and the influence of a

hot mantle plume (Ebinger and Sleep, 1998; George et al., 1998; MacDonald et al., 2001), rifting

in the WARS began essentially amagmatic with the earliest magmatism known to have occurred

closer to the cessation of broad regional extension (~48 Ma) (Wörner, 1999; Rocchi et al., 2002;

Ferraccioli et al., 2009; Granot et al., 2010). Extensive geochemical, isotopic, tectonophysical

and geochronological studies in the Northern Victoria Land region (e.g., Rocholl et al., 1995;

Salvini et al., 1997; Rocchi et al., 2002; Davey et al., 2006; Faccenna et al., 2008; Nardini et al.,

2009; Ferraccioli et al., 2009; Perinelli, 2011; Granot et al., 2010) have laid a robust framework

for constraining physical, temporal, and chemical influences on intraplate alkaline magmatism in

West Antarctica, including the Northwest Ross Sea (NWRS). This study assesses the

characteristics and complexity of magmatism across the transition from continental to oceanic

lithosphere in the NWRS where rifting occurred in the absence of a deep mantle plume. Previous workers have observed an increase in silica undersaturation coupled with increasing highly incompatible to moderately incompatible trace element ratios, increasing 206Pb/204Pb and

87Sr/86Sr ratios, and decreasing 143Nd/144Nd ratios along this transect (~164° to 174° East

longitude), suggesting differences in composition and degree of partial melting of source mantle

(Castillo et al., 2011). This study explores spatial variations in magmatic processes by comparing

petrography, mineral chemistry, T-P-H2O estimates, and oxygen isotopes from relatively undifferentiated alkaline volcanic rocks in the North West Ross Sea (NWRS; Figure 1) to 2

previously published and unpublished datasets (whole rock and mineral chemistry, Sr-Nd-Pb-

Os-O isotopes and 40Ar/39Ar dates) from Northern Victoria Land (NVL).

This study represents the first detailed examination of basalts collected along a transect that encompasses coincident continental and oceanic alkaline magmatism within the WARS. The

Cameroon Line in Africa is the only other known alkaline province on that straddles both oceanic and continental lithosphere with coincident volcanism. Studies involving the Cameroon

Line have yielded important contributions to our understanding of intraplate alkaline magmatism with respect to magma genesis, influences of plumes on melting, and tectonic controls on volcanism (Halliday et al., 1988; Rankenburg et al., 2005; Lee et al., 1996; Nkouatho et al.,

2008). Alkaline magmas are often attributed to low degree melting of enriched mantle source linked to recycling of oceanic lithosphere by mantle plumes (Hart, 1988; Weaver, 1991; Cauvel et al., 1992). The influence of mantle plumes generally produce large volumes of magmatism coincident (and/or prior) to extension and continental break-up. It has also been suggested that continental and oceanic alkaline compositions (basanite and alkali olivine basalt) are produced

by partial melting of metasomatic veins within the lithosphere (Pilet et al., 2005; 2008; 2010;

Chazot et al., 1997; Raffone et al., 2009). Since rifting in the WARS began amagmatic, and

pronounced magmatism did not begin until the end of extension, most authors favor the later

mechanism (Rocchi et al., 2002; Coltorti et al., 2004; Nardini et al., 2009; Perinelli et al., 2011;

Castillo et al., 2011).

The edge of rifted continents encompass a spectrum from “magma-dominated” to

“magma-poor” margins depending on whether magmatism is greater or less than expected from

the degree of lithospheric thinning and passive asthenospheric upwelling at normal mantle

temperatures (Sawyer et al., 2007). Magma-dominated margins are commonly associated with 3

elevated mantle temperatures related to a rising mantle plume and characterized by large-volume

igneous provinces (i.e. flood basalts) that erupted over short periods prior to or concomitant with

major extension (i.e. the East African Rift System). At magma-poor margins, like the NWRS,

continental breakup is tectonically driven, where continental lithosphere is extended and thinned

with little to no related magmatism. Lower volumes of magmatism may occur at magma-poor

rift margins from local mantle upwelling and decompression melting related to crustal thinning.

The transition from continental to oceanic lithosphere at magma-poor margins is believed to

have occurred over multiple rift phases by a process of strain localization and encompasses a

broad domain commonly referred to as the ocean-continent transition zone with characteristics similar to both oceanic and continental lithosphere (Sawyer et al., 2007; Péron-Pinvidic and

Manatschal, 2009). Rift margins are divided into a proximal, distal and oceanic domain based on the architecture and characteristics of the lithosphere (Péron-Pinvidic and Manatschal; 2009).

The proximal domain is characterized by stretched continental crust and half-graben structures related to early rifting. In the distal domain (which includes the ocean-continent transition), the crust is dramatically thinned (~15 to < 1 km) during subsequent rift phases. A zone of crustal necking and exhumed lithospheric mantle occur in the transition zone and has been explained by mid-crustal decollement and basement detachment faults (Péron-Pinvidic and Manatschal, 2009).

The transition from continental to oceanic crust is in most cases gradational and boundaries between domains are approximated at best.

In the NWRS, Cenozoic magmatism begins near the end of extension in the Ross Sea, where crustal lithosphere has been extensively thinned but no exhumed mantle has been documented. The emplacement of magmas have been closely linked with active lithospheric structures (Ferraccioli et al., 2009; Granot et al., 2010) where episodic melting exploited 4 metasomatically veined lithospheric mantle formed during an early amagmatic phase of extension (Rocchi et al., 2002). This study applies the concept of across-rift changes in the lithosphere to a land to sea geochemical and petrological traverse in order to investigate alkaline magmatism along a magma-poor rift margin. From here on I refer to the proximal domain as the continental rift shoulder (CRS) and the distal domain as the ocean-continent transition zone

(OCTZ).

The rift-axis of the WARS in the NWRS is roughly oriented NW-SE, therefore, sampling was designed to evaluate compositional differences as a function of perpendicular distance from the rift-shoulder (Figure 1), and divided among three domains based on the type of lithosphere associated with magmatism: 1) The CRS represents magmatism through continental lithosphere and includes samples from the Malta Plateau and inland from the Daniell Peninsula); 2) the

OCTZ represents magmatism through thinning, transitional lithosphere and includes basalts from the Adare Peninsula, across the continental shelf to the base of the continental slope (~1600 mbsl) in the Adare Basin; and 3) the oceanic domain which represents magmatism through what is considered to be true oceanic lithosphere and includes basalts from seamounts in the Adare

Trough, as well as the Balleny (~163° E, 67° S) and Scott islands (~ 180° E, 68° S). Other samples collected from the Hallet Peninsula, Daniell Peninsula and Melbourne Local suite are also considered to be within highly extended continental lithosphere but because they are located further south within the rift (Figure 1) they are not included within the continent to ocean transect.

The primary goals of this study are to: 1) Evaluate variability in magma composition across the continent-ocean transition, and 2) Differentiate between source and crustal influences on compositional variability. The discovery of varying degrees of textural, compositional and 5

oxygen isotopic disequilibrium in basalts collected from the OCTZ has elucidated a necessity to

explain the complex nature of magmatism in the transition from continental to oceanic

lithosphere. The broader impact of this study is to provide new insights on rifting and alkaline magmatism in the absence of a mantle plume for WARS and other magma-poor rifted margins worldwide.

6

2. GEOLOGIC BACKGROUND

The WARS is a broad region of extended continental lithosphere, stretching from the

northern Ross Sea to the Antarctic Peninsula. The division between thick continental craton and

thinned continental lithosphere of the WARS in the Ross Sea is bounded to the west by the

Transantarctic Mountains; an uplifted rift shoulder that extends more than 3500 km across the

continent (Fitzgerald et al., 1986; Stern and Uri, 1989; Stump and Fitzgerald, 1992). Crustal thicknesses decrease dramatically from 40-45 km beneath the Transantarctic Mountains to 20-25 km beneath the Ross Sea (Behrendt, 1999; Lawrence et al., 2006). Within the Ross Sea region are a series of north-south trending horst and graben structures (a.k.a. basins and highs) which

formed during the late Cretaceous and early Cenozoic (Cooper and Davey, 1987). The

topography at the base of the lithosphere which runs parallel to the rift axis has been identified

by seismic surveys and supported through theoretical modeling as a thin, narrow zone of

lithospheric thinning (Trey et al., 1999; Lawrence et al., 2006; Huerta and Harry, 2007). This

lithospheric necking zone is believed to have formed around the time that the WARS evolved

from broad regional extension to localized deformation and extension within the Ross Sea during

the late Paleogene (Huerta and Harry, 2007).

Models for Cenozoic magmatism in the WARS have shifted over the past decade away

from active plumes (Lanyon et al., 1993; Weaver et al., 1994; Storey et al. 1999) and fossil

plumes (Rocholl et al. 1995; Hart et al., 1997; Panter et al., 2000) to melting promoted by the

pre-existing metasomatic enrichment in the source (Rocchi et al., 2002; Finn et al., 2005; Nardini et al., 2009). Rocchi et al. (2002) and Nardini et al. (2009) proposed melting of a metasomatically enriched lithospheric mantle for NVL magmatism. According to their model, metasomatic enrichment of the mantle occurred during the late Cretaceous amagmatic rift phase. 7

Magmatism in the Cenozoic was triggered by craton-ward mantle flow in the Eocene-Oligocene and the reactivation of pre-existing trans-lithospheric faults. Faccenna et al. (2008) illustrates the link between tectonics and magmatism by integrating surface data and mantle tomography to reveal two deformational events following the middle Miocene: an early trans-tensional event coupled with sea-floor spreading that initiated magmatism (~40 Ma), and a subsequent extensional phase driven by mantle upwelling which lead to fault and fracture networks that produced the pathways for post-Miocene volcanism.

Cenozoic (~50 Ma to recent) basalts in the WARS are characterized by geochemical compositions similar to ocean island basalts (OIB), in particular those that have enriched radiogenic Pb isotopic signatures (e.g., 206Pb/204Pb > 19.5), indicating that they were derived from a HIMU-like (high time-integrated 238U/204Pb) mantle source (Rocholl et al., 1995; Hart et al., 1997; Panter et al., 2000; Sims et al., 2008; Nardini et al., 2009). Widely dispersed continental alkaline magmatism occurs throughout the southwest Pacific and is also characterized as HIMU-like (Finn et al., 2005; Panter et al., 2006). Additionally, the common occurrence of calcic amphibole in mafic alkaline magmas and xenoliths throughout the southwest Pacific (Johnson, 1989 and references there in; Gamble et al., 1987; Gamble et al.,

1988; Kyle, 1981; Coltorti et al., 2004; Cooper et al., 2007) is an indication that hydrous conditions were important in magma genesis. The broad distribution of alkaline magmatism with comparable chemistries suggests similarities in the magma source and the petrogenetic conditions. Recently discovered volcanic seamounts in the oceanic Adare Basin (Figure 1),

NWRS, were dredged in 2006/2007 during the NBP0701 marine geophysical survey cruise and preliminary geochemical analyses and chronological data indicate that the volcanism in the 8

Adare Basin is related to alkaline magmatism in the WARS (Panter, 2007; Panter and Castillo,

2008).

Variations in whole rock geochemistry with longitude have been reported for NWRS basalts (> 5 wt% MgO) and show and increase in ne, Nb/Y and La/Yb and decrease in TiO2/P2O5

with increasing longitude east (ocean-ward) from the CRS (Castillo et al., 2011). In general, it

has been proposed that magmas become more alkaline from the CRS to ocean as degree of

partial melting decreases. Likewise, it has been shown that magmas become less affected by

contamination from continental lithosphere as evidenced by decreasing 87Sr/86Sr and 187Os/188Os

and increasing 143Nd/144Nd (Castillo et al., 2011; Juda et al., 2012).

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3. METHODS

3.1. Sample Summary

A total of 21 mafic and one mafic dike were selected from throughout the NWRS to

represent continental, oceanic, and continental-oceanic transitional magmatism (Table 1). An

additional 4 slightly to moderately evolved samples were selected to assess differentiation. The

samples include newly dredged seamounts from the Adare Basin (Panter and Castillo, NBP0701)

along with samples from various NVL collections (Polar Rock Repository at Byrd Polar

Research Center, Ohio State University; Stan Hart at Woods Hole Oceanographic Institution;

Philip Kyle at New Mexico Tech; and Nick Mortimer at GNS Science, New Zealand). New

mineral chemistry and oxygen isotopes from this study expand the existing published (Hamilton,

1972; Johnson et al., 1982; Muller et al, 1991; Rocholl et al, 1995; Horning and Wörner, 2003;

Mortimer et al., 2007; Panter, 2007; Panter and Castillo, 2008; Castillo et al., 2011; Juda et al.,

2012) and unpublished (S. Hart & P. Kyle, K. Panter & P. Castillo) datasets (whole rock major

and trace elements, Sr-Nd-Pb-Os isotopes, and dating by 40Ar/39Ar or K-Ar methods) for these

samples.

Care was taken to select fresh basaltic lavas that were relatively undifferentiated (MgO ≥

6 wt%; Mg# > 50; Ni = 70 – 205 ppm; Cr = 135 – 540 ppm), to minimize effects of

differentiation and possible contamination by crust during magma ascent to the surface. The

most primitive samples are basanite and alkaline basalt with differentiation occurring along the

basalt-trachyte and basanite- magma series (Figure 2). Samples from the Melbourne

Volcanic Province (Figure 1; S. Hart, P. Kyle, and G. Wörner) represent magmatism through

continental lithosphere (e.g. CRS) and are characterized by > 21% hypersthene normative (hy) to

< 3% nepheline normative (ne) alkaline olivine basalts (Mg# 56 – 59; Ni = 86 – 93 ppm; Cr = 10

272 – 349 ppm). Oceanic samples including Adare Basin Seamounts (ABS) and Balleny Island are mostly basanite (Mg# 53 – 70; Ni: 106 – 338 ppm; Cr: 196 – 540 ppm), with one evolved sample, tephriphonolite, from Scott Island (Mg# 10; Ni and Cr contents < 10 ppm; P. Castillo &

K. Panter, S. Hart & P. Kyle unpublished data). The OCTZ includes ABS from the base of the continental slope (~1600 mbsl), , Adare Peninsula, Hallet Peninsula, Daniell

Peninsula, and Melbourne local suite (Figure 1). Mafic lavas are a mix of alkaline olivine basalt, hawaiite, basanite, and (Mg# 47 – 68; Ni = 71 – 205 ppm; Cr = 135 – 477 ppm), with evolved tephrite and mugaerite (Mg# 37 – 43; Ni < 40 ppm; Cr < 70 ppm). One intrusive, a basaltic dike from the Adare Peninsula (P74833; Mg# 56; Ni = 99 ppm; Cr = 304 ppm) was included because of its comparable age and composition to CRS volcanism and presence of amphibole. The four evolved samples (e.g., Mg# 10 – 43; A214B, SC-002, PRR-3873 and SCI-

002) were selected to increase the number of samples containing amphibole and to exemplify magmas that have been significantly affected by differentiation.

3.2. Sample Preparation

Rocks were crushed and sieved in order to separate between 250 μm and 2 mm in diameter. Approximately 4-10 mg of the best olivine, clinopyroxene, and amphibole mineral separates (i.e. pristine cores, few to no inclusions and no alteration) were chosen for oxygen using laser fluorination (LF). Four additional olivine grains from each sample were set aside for oxygen analysis by secondary ion (SIMS). Samples chosen for analysis by SIMS were either whole phenocrysts or fragments (either fresh or with minimal alteration at rims) and contain varying percentages of melt and oxide inclusions which 11

(unlike LF) are easily avoided by the SIMS laser (~10 µm diameter spot size with ± 0.3 ‰ precision). From the remaining olivine, amphibole, clinopyroxene, and plagioclase separates, 4 to 8 of the next best grains (whole grains when possible, and are representative of the separates analyzed for oxygen) were set aside to be mounted for electron microprobe analysis.

Olivine grains for SIMS were divided among three 2.5 cm mounts based on relative size.

Grains were placed within 5mm from the center of each mount to minimize grain topography from polishing (Kita et al, 2009). San Carlos olivine standards were placed at the center of each mount. High precision polishing was performed on grain mounts at UW-Madison to minimize surface relief to less than 1 micron (Kita et al., 2009; Valley and Kita, 2009; Eiler et al., 2011).

Reflected light and back scattered electron (BSE) images were taken of each grain prior to analysis to assess compositional zoning and select locations on individual grains to be measured in situ by microprobe and SIMS.

3.3. Mineral Chemistry

Six to eight phenocrysts from each sample separate (olivine, clinopyroxene, amphibole and plagioclase) were mounted and singly polished on 1 cm diameter epoxy rounds. These were then secured into 1 inch brass mounts (three rounds per mount) and coated for analysis by electron microprobe. Major and minor element chemistry was measured on a Cameca SX-100

Electron Microprobe Analyzer at the University of Michigan EMAL laboratory using a 15kV accelerating voltage beam, 15nA beam current, and 10-12μm spot size. A total of 307 unknowns were analyzed; 174 olivine, 76 clinopyroxene, 23 amphibole, and 38 plagioclase grains. The following elemental routines were performed for each mineral: olivine = Si, Cr, Fe, Mn, Mg, Ca, 12

P, Ni, clinopyroxene =Si, Al, Ti, Cr, Fe, Mn, Mg, Ca, Na, K, P, Ni, Cl, plagioclase = Mg, Al, K,

Ca, Ti, Cr, Mn, Fe, Na, P, Si, Sr, Ba, and amphibole =Mg, Al, K, Ca, Ti, Cr, Mn, Fe, Na, P, Si,

Cl, F. Measurements, a total of 475 unknowns, were taken primarily at mineral cores (or

presumed cores on fragments), and also at rim and intermediate grain locations when

compositional or textural variations were observed in BSE or thin section. From here on, the

number of unknowns analyzed - na and number of rock samples - ns will be used in the description of the data.

Detailed microprobe transects were analyzed on six olivine grains that exhibit compositional zoning (MA-009a d, A227B d, PRR-5171 c, PRR-3872 a, A232B d, and D12-1 c).

The purpose of these transects was to investigate chemical zoning with respect to magmatic processes (i.e. differentiation and magma mixing/mingling) and will be used in future work to assess temporal constraints on magmatic processes among across-rift domains (cf. Costa et al.,

2008). Olivine grains were selected for detailed transects based on the following criteria: 1) compositional zoning that exhibit 2 to > 5 mol% Fo variation between core and rim, 2) whole, or mostly whole, olivine crystals with clear grain boundaries, preferably with groundmass still attached, and 3) grains from samples that are broadly distributed across continental and oceanic domains. Transects vary from 47 to 322 μm with 1-5 µm steps in measurements depending on grain size and extent of zoning.

3.4. Oxygen Isotopes

Minerals analyzed for oxygen isotopes were measured in 1.6 to 2.6 mg aliquots of olivine

(ns = 10), clinopyroxene (ns = 17), and amphibole (ns = 5) by laser fluorination (LF) at the 13

University of Wisconsin-Madison which is equipped with a CO2 laser probe system (BrF5

reagent) attached to a dual-inlet five-collector Finnigan/MAT 251 mass-spectrometer. In cases

where a given sample yielded two size populations of phenocrysts, aliquots were lumped

comparatively by size. Measurements for unknowns were corrected by the average difference

18 between measured and accepted values for UWG-2 Gore Mountain garnet standard (δ OVSMOW

= 5.80 ‰) measured on the same day. Daily standard deviation for UWG-2 was 0.08 to 0.10 ‰

(2σ, na = 8). For consistency, the larger of these two is used to represent precision for all

unknown measurements. Duplicate measurements of unknowns had an average reproducibility

of ± 0.21 ‰ for olivine, ± 0.38 ‰ for clinopyroxene, and ±0.13 ‰ for amphibole (2σ).

Duplicates were attempted for all runs, but were only obtained on 60% due to elimination by

more rigorous inspection of grains prior to analysis and also occasional failed analysis.

In-situ oxygen isotope compositions were measured for an additional 96 olivine

phenocrysts (ns = 25, ~4 grains/sample) by SIMS at the University of Wisconsin-Madison

WiscSIMS Laboratory. Oxygen isotope analysis by SIMS was based on two objectives: 1) to

obtain oxygen data on samples that were unsuitable for laser fluorination as described above, and

2) to assess intra-crystalline and intra-sample heterogeneities (both compositionally and in

oxygen isotopic ratios). A CAMECA IMS-1280 large-radius, double-focusing secondary-ion

mass-spectrometer was used with a primary beam intensity of 1.5 to 1.8 nA (spot size ~10 μm).

Since olivine compositions range Fo60 to Fo91, an Mg# (Fo) calibration curve (Valley and Kita,

2009) was used to correct for bias in oxygen values relative to San Carlos olivine (SCOL) standard. A bracketed average of SCOL was run approximately every 15 analyses, with a standard precision of ± 0.3 ‰ (2σ). Eight unknowns representing the most extreme (lowest and

18 highest) δ Ool values were selected (following initial inspection of oxygen analysis by SIMS) 14 and an additional 5 to 6 analyses were taken at cores to test intra-grain reproducibility; from here on these will be referred to as Tavg. SIMS averages are typically 0.1 to 0.2 ‰ heavier than LF averages. SIMS measurements were taken primarily at olivine cores, which may have produced a sampling bias contributing to the systematic difference between SIMS and LF values.

15

4. RESULTS

4.1. Petrography

Petrography of thin sections was primarily used to assess paragenesis and mineral-melt equilibrium. Note that six of the 26 samples do not have representative thin sections due to the nature of some rock samples (fragments only; Table 2). For these samples, fragments were mounted on slides and polished to 30 – 50 µm. Overall, samples are porphyritic with either holocrystalline pilotaxitic or trachytic groundmasses or holohayline to hypohayline groundmasses, which are most prominent in seamount samples (Table 2). Olivine and clinopyroxene ± plagioclase and rare amphibole are the most common phases.

Plagioclase and are common groundmass phases with subordinate clinopyroxene, olivine and amphibole. In a few samples plagioclase is absent altogether.

Olivine is euhedral to subhedral and often contains inclusions of Fe-Ti oxides up to 2 vol

%. Clinopyroxene is euhedral to subhedral with up to 3 vol% Fe-Ti oxide inclusions. Larger phenocrysts of clinopyroxene (≥ 0.5 mm) often exhibit disequilibrium textures with spongy cores

(Figure 4d) and prominent zoning, while smaller phenocrysts (< 0.5 mm) are euhedral to subhedral and are only occasionally zoned. Pinkish-brown glass is observed in large, embayed clinopyroxene phenocryst (~1 mm) from the Possession islands (PRR-3842 and PRR-3873;

Figure 3c). Glomerocrysts include euhedral to anhedral olivine and clinopyroxene crystals are often found in samples from the Adare Peninsula and the Possession islands and range from 1 to

2 mm in diameter (Figure 3a and b). In some cases the clinopyroxene is deformed (kink banded), suggesting that some of these crystal clots may be xenoliths (Figure 3d). Ultramafic cumulates, crustal and mantle xenoliths are common in basalts from NVL and the rest of the

WARS (Nardini et al., 2009; Armienti and Perinelli, 2010; Perinelli et al., 2006, 2011; Berg, 16

1991; Wörner, 1999; Gamble and Kyle, 1987; Coltorti et al., 2004). For this reason, larger (> 1

mm) clinopyroxene grains were generally avoided during hand-picking of NWRS mineral

separates, but in many cases fragments of grains were included in analysis of mineral chemistry

(see discussion on mineral-liquid equilibrium in section 4.4). Plagioclase is less common as a

phenocryst phase, present in less than half of NWRS basalts. Plagioclase is typically euhedral

and zoned, though sometimes subhedral to anhedral with sieve texture and reaction rims (Figure

4a and b). Rare amphibole phenocrysts commonly exhibit mild to extensive reaction rims with

fine grained crystallization of opaque oxides and other silicate mineral phases (Figure 4c).

Groundmasses contain varying amounts of plagioclase in most samples but it is rare to absent in seamount samples. Also in seamount samples, quenched textures such as variolitic olivine and clinopyroxene and skeletal crystals of olivine are common (Figure 5a-c), whereas samples collected in the CRS and OTCZ often show disequilibrium textures that consist of spongy and

zoned clinopyroxene, reaction rims on amphibole, and sieved plagioclase.

4.2. Mineral Chemistry

The composition of olivine phenocrysts from the NWRS are highly variable with core values

ranging from Fo66 to Fo91 (Figure 6 and Table 3). Most olivine display weak to strong normal zoning (1-17 mol% decrease in Fo), with rare reverse (up to 3 mol% increase in Fo) and oscillatory zoning (in OCTZ basalts only). The highest Fo olivine (Fo ≥ 89) are found in the oceanic (D4-1, ABS) and OCTZ (A210B Adare Peninsula) domains. Olivine from seamounts generally have higher Fo (79-91) and Ni contents (732- 3546 ppm) and weaker zoning (<3% Fo from core to rim) than from Balleny Island (phenocryst: Fo 71 to Fo88, Ni contents from 86 to

2923 ppm, and strong normal zoning up to 14 mol% Fo; xenocrysts: Fo75 and Ni contents from 17

1275-1536 ppm). In general, Ni decreases and Mn increases in olivine with decreasing Fo and

whole rock Mg# as expected with differentiation by fractional crystallization of olivine ±

clinopyroxene, magnetite (Figure 7). varies between 450 and 3164 ppm and shows no

correlation to Fo or whole rock Mg #. The highest Ca is found in basalts from the OCTZ and

lowest Mn from the CRS. Overall the composition of olivine in basalt is more variable than

olivine measured in ultramafic cumulates (Fo73 to Fo84; Ni: 629-3379 ppm; Mn: 1007-3330 ppm;

Ca: 715-3359 ppm) and mantle xenoliths (Fo84 to Fo91; Ni: 786-5894 ppm; Mn: 77-2711 ppm;

Ca: 357-1858 ppm) from this region of NVL (Perinelli et al., 2011; 2006).

Detailed microprobe transects were collected on six olivine that exhibit strong (ΔFo ≥ 5

mol%), oscillatory or abrupt zoning from various locations across the continent to ocean transect

(Malta Plateau, Adare Peninsula, Possession Islands and ABS). Core to rim transects generally

show a decrease in Fo and Ni, and an increase in Ca and Mn from core to rim (Figure 7a-f).

Oscillatory zoning is observed in samples A227B and PRR-3872 from the OCTZ (Figure 7b and

7e). For olivine in sample A227B, initial normal zoning is halted by abrupt reverse zoning and increased Ca and Ni, followed by normal zoning and decreased Ca and Ni (Figure 7b). For

olivine in sample PRR-3872, a homogeneous low Fo core is followed by gradual reverse zoning

with increased Ca, followed by abrupt normal zoning with decreased Ni and increased Ca

(Figure 7e). The other grains have homogeneous cores with normal zoning occurring either

gradually (PRR-5171_c and MA-009a_d) or abruptly (A232B_d and D12-1c) at rims (Figure 7).

Clinopyroxene phenocrysts are diopside and contain high calcium contents with samples

from oceanic domains plotting above the 50% Ca line on the quadrilateral shown in

Figure 6 (see also Table 4). Clinopyroxene in samples from the CRS plot below and from the

OCTZ plot both above and below the 50% Ca line (Figure 6) . Clinopyroxene cores range in 18

composition (Wo37-52, En31-52, Fs7-21 and Mg#cpx 60-90) and display weak to strong normal zoning

(as much as 3.3 % decrease in Mg#cpx from core to rim) and less common reverse zoning (up to

3.3 % increase in Mg# from core to rim; OCTZ only). Clinopyroxene in NWRS basalts have Ti,

Na, Cr and Ca concentrations that are comparable to clinopyroxene in xenoliths from NVL

(Perinelli et al., 2006) and Foster Crater in the southern Ross Sea (Gamble and Kyle, 1987) but have higher Fe, Mn, and Mg contents. In general, Ni decreases while Ti and Mn increase with decreasing Mg#cpx (expected for differentiation by fractional crystallization). Core-rim variations in Al, Na, Cr and Ca concentration do not behave consistently with changes in Mg#.

Phenocrysts of amphibole are present in eight out of 25 NWRS basalts, mostly occurring in whole rock compositions with lower Mg# (< 56) and Ni contents (< 160 ppm). Most of the amphibole is classified as kaersutite (Leake et al., 1997) with Mg#amp 59 – 72. Exceptions are the magnesiohastingsite in P74833 (dike) and hastingsite to ferropargasite in tephriphonolite sample

SCI-002 (Figure 6 and Table 5). Mg#amp varies as much as 10 mol% within a sample, and observed core-rim zoning is reverse (~1-3 mol% Mg#). Kaersutite is common in basalts from other areas within the WARS, including ultramafic cumulates, megacrysts and xenoliths

(Gamble and Kyle, 1987; Rocchi et al., 2002; Coltorti et al., 2004; Perinelli et al., 2006; 2011;

Cooper et al., 2007).

Non-groundmass plagioclase is found in less than half of NWRS basalts and has core compositions that range from An40 to An70. The majority of these plagioclase are considered to be antecrysts (i.e. crystals that formed earlier in the differentiation process and have been reincorporated into the magma) because they exhibit sieve texture, reverse zoning (2 to 22 mol% increase in An from core to rim; Table 6), and in rare cases xenocrysts when they exhibit deformation twins (i.e. Balleny Island; Figure 4a). Only a few samples contain phenocryst 19

plagioclase (SCI-002, PRR-3873, and A214B); in all cases whole rock compositions are more

evolved (Mg# < 43).

4.3. Oxygen Isotopes

Oxygen isotope analyses for all minerals are averaged and reported at 2σ standard

deviation (Table 7). Olivine from Balleny (24583) was divided into 2 populations (phenocryst

and xenocryst) due to the observation of xenolithic clots in hand sample and thin section.

Phenocryst olivine was characterized by its smaller diameter (~0.5 mm), euhedral texture and

higher average Fo in cores (79 ± 18 mol%). Xenocrysts are characterized by larger diameter

grains (> 1 mm), subhedral to anhedral textures and lower Fo in cores (75 mol%). Oxygen

isotopes for clinopyroxene and amphibole measured by LF are compared to olivine δ18O measured by both LF and SIMS (sample average) methods with the purpose of evaluating oxygen isotope fractionation (Δ18O) between coexisting minerals.

4.3.1. Oxygen isotopes by laser fluorination

NWRS olivine measured by LF average 5.01 ± 0.18 ‰ (2σ, ns = 10, na = 15) and range

from 4.69 to 5.30 ‰ (Figure 8a). Clinopyroxene average 4.97 ± 0.36 ‰ (ns = 17, na = 26) and

range from 4.38 to 5.89 ‰. Amphibole average 5.17 ± 0.10 ‰ (ns = 6, na = 11) and range from

4.96 to 5.30 ‰. Overall, the average LF values for olivine in this study are slightly lower than

18 average LF values reported for other NVL basalts (δ Ool = 5.29 ± 0.40‰, 2σ; Nardini et al.,

18 18 2009) and ultramafic cumulates (δ Ool = 5.29 ± 0.46‰, 2σ; Perinelli et al., 2011). The δ O values of clinopyroxene are also generally lower than those measured in NVL ultramafic cumulates (5.28 ± 0.29 ‰, 2σ; Perinelli et al., 2011). Clinopyroxene and olivine in NWRS 20

18 samples also show overall lower Δ Ocpx-ol values (discussed below) relative to NVL ultramafic

18 cumulates and mantle xenoliths (Figure 9). Large standard deviation in δ Ocpx from duplicate LF

analyses suggests intra-crystalline heterogeneities in clinopyroxene, either among or within

18 grains. Unfortunately, intra-crystalline variability in δ Ocpx was not tested with SIMS since the

analytical uncertainties for in-situ clinopyroxene analysis are not as well constrained as they are

for olivine.

4.3.2. Oxygen isotopes by SIMS

NWRS olivine measured by SIMS, including core and rim analyses, range between 4.50

18 to 5.78 ‰ and average 5.15 ± 0.5‰ (2σ, ns = 23, na = 149; Figure 8a). In-situ variations in δ Ool

are as much as 0.56 ‰ lighter to 0.50 ‰ heavier at rims than cores (Figure 8b). However, this is

not statistically significant with respect to instrument precision (± 0.28 ‰ at 2σ) and therefore

any intra-crystalline variability that may exist is unresolvable by SIMS. The lowest single spot

18 δ Ool values are from samples PRR-5169 and PRR-5171 (4.50 ‰) with sample average values being slightly higher 4.82 ± 0.36 ‰ (na = 7) and 4.71 ± 0.30 ‰ (na = 7), respectively. The

18 highest single spot δ Ool values are from MA-009a (5.76 ‰) and MA-117 (5.78 ‰) with sample

averages of 5.44 ± 0.43 ‰ (na = 8) and 5.39 ± 0.56 ‰ (na = 7), respectively. Sample 24583

phenocrysts and xenocrysts average 5.19 ‰ ± 0.40 and 4.94 ‰ ± 0.33 (SIMS), respectively. The

18 difference between single spot analysis of high and low δ Ool values are resolvable beyond

analytical precision. Core averages (Tavg) from individual olivine with the highest and lowest

18 δ O have standard deviations ~20-60% less than (i.e. 0.05 to 0.43 ‰; 2σ) sample averages

which average multiple grains (Figure 8a). 21

Fractionation of δ18O between mineral phases is in most cases non-equilibrium for clinopyroxene-olivine, and generally positive for amphibole-olivine (Figure 9a and b). There is a

18 wide range in Δ Ocpx-ol from +0.96 to -0.67 ‰ (Table 7; Figure 9), where the expected value for

18 Δ Ocpx-ol in equilibrium with a melt at 1200 °C (~mantle) is +0.4 ‰ (Mattey et al., 1994; Eiler,

2001). Fractionation between amphibole-olivine range tightly from -0.12 to +0.33 ‰ (Table 7;

Figure 9b). Equilibrium fractionation between olivine and amphibole is not as well constrained,

18 however Chazot et al (1997) reported Δ Oam-ol ~ +0.2 ‰ for a hydrous spinel lherzolite, which

18 agrees with the findings in this study. It is unlikely that the variation in Δ Ocpx-ol is related to

differentiation because of the relatively unfractionated compositions, and because clinopyroxene

is in many cases isotopically lighter than olivine (i.e. differentiation by fractionation

crystallization of mafic minerals would cause δ18O values to increase). In addition, there is no

18 correlation between δ Ool and Fo content (Figure 8a), which would also be expected for differentiation. It is important to note, however, that olivine rims analyzed by SIMS are

generally higher in δ18O than cores for olivine with lower Fo contents (< 74%; Figure 8b). This

may explain why olivine analyzed by LF on multiple whole grains is typically 0.1 to 0.2 ‰

lighter than SIMS values for the same sample, which is based predominately on average core values.

18 The difference in Δ Ocpx-ol values that is displayed in Figure 9 could possibly reflect a

range of crystallization temperatures. However this cannot explain negative values. Since LF

18 requires bulk mineral separates, δ Ocpx represents an average for multiple grains, and therefore

may include both minerals in equilibrium and disequilibrium with olivine. It may also be biased

in the presence of significant amounts of oxide mineral inclusions (e.g. magnetite and chromite),

16 18 18 which are typically enriched in O with respect to O. For example at 1200 °C the Δ Omt-ol is 22

18 equal to -1.4 ‰ compared with the Δ Ocpx-ol which is equal to +0.4 ‰ (Eiler, 2001). Since

many clinopyroxene grains contain 1-3 vol% Fe-Ti oxide inclusions, I performed a simplified

mass balance calculation to assess the effect of Fe-Ti inclusions on the δ18O of clinopyroxene

18 18 measured by LF. The mass balance was calculated assuming δ Ool = 5.00 ‰, Δ Ocpx-ol = +0.40

18 ‰ and Δ Omt-ol = -1.40 ‰ at 1200 °C. The calculations indicate that it would take ~20 vol%

magnetite to reduce the δ18O of clinopyroxene from 5.40 to 5.00 ‰. To explain the maximum

18 observed difference between olivine and clinopyroxene (Δ Ocpx-ol = -0.96 ‰) by magnetite

‘contamination’ would require an inclusion abundance of > 50 vol%. The maximum volume of

magnetite inclusions observed in clinopyroxene is ≤ 5% and, therefore, it is unlikely that

18 extremely low Δ Ocpx-ol values (< 0) can be explained by ‘contamination’ but are most likely attributed to magmatic disequilibrium between minerals and melt.

4.4. Thermobarometry and water estimates

Olivine, clinopyroxene and amphibole chemistry was used to estimate equilibrium

temperatures and pressures for NWRS samples shown in Table 8 using a variety of hydrous and

anhydrous thermobarometric equations (Beattie, 1993; Putirka et al., 2007; 2003; Putirka, 2008;

Ridolfi et al., 2010). For comparison, amphibole T-P-H2O was also estimated for NVL cumulate

mineral chemistry reported by Perinelli et al. (2011). Whole rock chemistry was used to

represent the liquid composition for mineral-liquid thermobarometers. KD(Fe-Mg) for olivine- liquid equilibrium exchange is 0.30 ± 0.06 (2σ) based on experimental data (Roeder and Emslie,

1970). KD(Fe-Mg) for clinopyroxene-liquid equilibrium exchange is ~0.27 ± 0.06 (2σ) based on experimental data (Putirka, 2008). Both cores and rims of minerals were tested for Fe-Mg equilibrium exchange (Figure 10). In general cores were favored for T-P calculations, however 23

rims were included in a few cases where: 1) equilibrium KD(Fe-Mg) values were observed, and

2) differences between T-P estimates from equilibrium cores and rims were <25°C and <0.3

GPa. However, caution must be taken when interpreting T-P based on KD (Fe-Mg) mineral-liquid equilibrium values since some samples contain glomerocrysts and xenoliths which may have biased the concentrations of whole rock Fe and Mg. For this reason, petrography was used in concert when assessing the nature of disequilibrium on a case-by-case basis.

Nearly half of all olivine cores plot below the equilibrium range, though many overlap the equilibrium field with rims trending above equilibrium (Figure 10a). Gee and Sack (1988)

ol-liq have shown that KD (Fe-Mg) decreases with decreasing SiO2 or increasing alkalis, ranging between 0.33 and 0.27 from tholeiite to alkali basalt, however the generally accepted Roeder and

Emslie (1970) model of ~0.30 and a 2σ standard deviation (± 0.06) encompasses this variability.

Note a few samples lie almost entirely above (e.g. P74833), below (e.g. PRR-3873), or bimodally above and below the equilibrium range (e.g. MA-117 and 24583). Values above equilibrium may result from differentiation or mixing processes. The Rhodes diagram (Figure

11a) depicts differentiation and olivine removal as controlling processes for disequilibrium, although mixing cannot be excluded. Cores for clinopyroxene are either at or above equilibrium with the whole rock, while rims are generally above the equilibrium range but overlap with core values (Figure 10b). Clinopyroxene-liquid Na-Al and Ca-Na equilibrium exchange was also tested by comparing observed and predicted pyroxene end-members (i.e. DiHd, EnFs, and CaTs;

Figure 11b). Data generally plots on or above the 1:1 line in Figure 11b, which is in agreement

cpx-liq with results for KD(Fe-Mg) .

Comparison of KD(Fe-Mg) for olivine-liquid and clinopyroxene-liquid show

inconsistencies between the two phases and whole rock composition (Figure 10). Eight of the 26 24

samples have clinopyroxene cores well above equilibrium (> 0.4), while coexisting olivine is

essentially in equilibrium with the whole rock. Balleny Island (24583) has no olivine or

clinopyroxene cores in equilibrium with the whole rock. Both olivine and clinopyroxene are consistently above equilibrium for P74833 (Adare Peninsula), and below equilibrium for PRR-

3873 (Possession Islands). Clinopyroxene-liquid equilibrium was also tested be comparing observed and predicted pyroxene end-members (i.e. DiHd, EnFs, and CaTs) which generally plot on or above the 1:1 line (Figure 11b).

Water was estimated by plagioclase-liquid hygrometers (Lange et al., 2009 and Putirka,

2008) and by amphibole thermobarometry (Ridolfi, 2010). Plagioclase-liquid hygrometers are conventionally used to estimate H2Omelt, but are limited by their sensitivity to mixed volatile

systems. Specifically, the hygrometer of Lange et al (2009) has the lowest error (± 0.32 wt%),

but only when strict conditions are met: 1) temperatures between 825° - 1100°C, 2) pressures

between 0.05 to 0.3 GPa, and 3) a pure H2O volatile system. Condition 1 is reasonable for this

study; however, pressures estimated from cpx-liq are generally higher than condition 2. While it

is expected that plagioclase crystalized at lower pressure/temperature than clinopyroxene, this is

the only reasonable method for T-P estimation without creating a circular argument (e.g. water

estimates are T sensitive). Condition 3 is the most difficult to satisfy since most magma systems

are not pure H2O volatile systems, and the effects of mixed volatiles on T-P and H2O estimates

are still poorly understood. For instance, Mt. Erebus, an active alkaline volcano in Southern

Victoria Land, is a high CO2 producer with up to 0.7 wt% reported in olivine hosted melt

inclusions from basanite (Eschenbacher et al., 1998). While the hygrometer of Putirka (2008)

offers fewer conditional restrictions, its error is significantly higher (± 1.1 wt%). For these

reasons, water estimated from plagioclase will act as a rough minimum for NWRS magmas. 25

Likewise, estimates from amphibole will serve as a rough maximum for comparison purposes

only.

4.4.1. Olivine thermometry

Olivine-liquid equilibrium temperatures range from 1127° to 1346 °C (ns = 22, na = 62)

using the anhydrous thermometer of Beattie (1993; Table 8). Temperatures estimated for

hydrous systems (Putirka et al., 2007) and CO2-bearing systems (Sisson and Grove, 1993) are

35° ± 34°C and 99° ± 19°C cooler than anhydrous temperatures (Beattie, 1993), respectively.

Since olivine is expected to crystallize at higher temperature than pyroxene under the same set of

conditions, the anhydrous thermometer of Beattie (1993) will be used from here on since it yields

more consistent temperature estimates near or above those of clinopyroxene.

4.4.2. Clinopyroxene thermobarometry

Clinopyroxene-liquid equilibrium temperatures and pressures range from 1102° to 1285

°C and 0.7 to 1.4 GPa (ns = 14, na = 51). Clinopyroxene-only temperatures and pressures range

from 1077° to 1279 °C and 0.3 to 1.2 GPa for samples that achieved cpx-liquid equilibrium (ns =

14, na = 51). T-P estimates from cpx-only methods tend to be lower than those from cpx-liq

methods, but generally within error (Table 8). Temperatures estimated from olivine are 40 °C

cooler to 40 °C warmer than those estimated from clinopyroxene, however, these differences are

within error of the thermometers (± 44° for olivine and ± 60° for clinopyroxene).

4.4.3. Amphibole thermobarometry

Kaersutite is a calcic amphibole with characteristically high Ti and ivAl contents and has been

shown experimentally to exist under higher P-T conditions and lower H2O concentrations than 26

other forms of amphibole (Ridolfi et al., 2010; Ridolfi and Renzulli, 2011). The sensitivity of

amphibole composition to T-P conditions is assessed through Ti- and Al- Tschermak

substitution. Al-Tschermak substitution (TSi → TAl and M1-M3Mg → M1-M3Al) in the amphibole structure is pressure sensitive, while Ti-Tschermak substitution (TSi → TAl and M1-M3Mn → M1-

M3Ti) is sensitive to temperature (Spear, 1981; Johnson and Rutherford, 1989; Bachmann and

Dungan, 2002). Amphibole from the NWRS exhibit strong temperature dependence (Figure 12a)

and considerable pressure dependence (Figure 12b) on compositional variability, and are

comparable to amphibole in ultramafic cumulates and mantle xenoliths from NVL and Foster

Crater (with the exception of amphibole from Mt. Melbourne and Mt Overlord; Coltorti et al.,

2004).

Amphibole T-P estimates are 956° to 1068 °C and 0.2 to 0.8 GPa (ns = 7, na = 22). Although

most samples plot just outside the range of thermal stability defined by Ridolfi et al. (2010)

shown in Figure 13, amphibole from NVL mantle xenoliths overlap the range of NWRS samples.

Again, the thermobarometer is best suited for pure H2Omelt systems, and the effects of mixed

volatiles on amphibole stability are not well constrained. However, experimental work of Fresie

et al. (2009) has shown that mixed-fluid systems (i.e. XH2O < 1) result in higher thermal stability

of calcic amphiboles than pure-H2O systems (i.e. XH2O = 1), where thermal stability is expected

2- - to increase with increasing CO3 and OH dissolved in the melt. Similarly, -substitution

has been shown to result in higher thermal stability of pargasite (another form of calcic amphibole found almost exclusively in alkaline magmas). Since there is good agreement in Ti- and Al- Tschermak substitution between amphibole in this study and other WARS xenoliths

(Figure 12), it is reasonable to assume these are high T-P amphibole and may suggest the

presence of mixed volatiles in the melt. 27

4.4.4. Water estimates

Water estimates are highly variable for NWRS basalts, especially when comparing methods. Estimates using the plag-liquid hygrometer of Putirka (2008) are extremely low and in many cases negative which is not realistic (Table 8). Water estimates using the plag-liquid hygrometer of Lange et al. (2009) range from 0.3 to 2.9 wt% (ns = 10, na = 36). Unfortunately, the lack in abundance of phenocryst plagioclase and the regular occurrence of sieve texture observed in nearly all samples makes these estimates highly suspect. H2Omelt estimated from all amphibole compositions are between 0.7 and 6.1 wt% (ns = 7, na = 22). Estimates for H2Omelt from kaersutite (sensu stricto) are 2.3-6.1 wt%, with H2Oamp from 1.7-1.9 wt%. These estimates are erroneously high for basaltic magmas, especially in a rifted-continent setting. Volatile measurements from olivine hosted melt inclusion in Mt. Erebus basanite from the same province are ~1.5 wt% H2O, and 0.6 wt% CO2, with F, S and Cl approximately 1800, 2100 and 800 ppm, respectively (Oppenheimer et al., 2011). Assuming that the volatile compositions are similar for other WARS basanite; this would suggest that the stability of kaersutite is controlled by mixed volatile compositions and not by water alone (the effects of which are not well constrained for amphibole stability). So while T-P estimates from amphibole are believable, estimates for

H2Omelt are not. This leaves us with no reasonable way to constrain primary volatile concentrations; however the presence of high T-P amphibole in NWRS basalts is an indicator of volatile-rich conditions.

28

5. DISCUSSION

5.1. Magma Origin

Magmas that are considered the least fractionated from primitive mantle melts will have

high whole rock Mg# (> 66 mol%), Ni (> 170 ppm) and Cr (> 400 ppm) contents (Green and

Ringwood, 1970) and low 87Sr/86Sr and 187Os/188Os, as well as high 143Nd/144Nd ratios. The

basalts should also have a low abundance of crystals with olivine as the dominant mineral phase.

Furthermore, if the composition of the mineral assemblage yields thermobarometric results

indicating high pressure and temperature conditions (i.e. Moho depth or greater) during

crystallization, then it seems reasonable to speculate that very little differentiation, and potential

interaction with the crust, occurred while the magmas rose to the surface. The low abundance of

phenocrysts, high whole rock Mg#, Ni and Cr contents and uncontaminated Sr, Nd, Pb and Os

isotopic signatures for most of the NWRS basalts analyzed in this study are optimal for evaluating early differentiation processes and can provide clues for mantle source compositions.

Differentiation of magmas in the NWRS was controlled, primarily, by olivine and clinopyroxene removal, which is indicated by the decrease in Ni and Cr contents with whole-rock Mg# shown in Figure 14a. The increase in Sr concentration with decreasing Mg# (Figure 14b) suggests that plagioclase was not being removed.

Olivine chemistry can provide information about near-primary magma conditions since it is the first phase to crystallize from the melt. In the case of NWRS basalt, only olivine cores with high Fo (≥ 85) and high Ni contents (≥ 1200 ppm) are considered to record a minimal amount of crystallization subsequent to extraction from the mantle. Olivine trace elements (Ni and Mn) were used to assess the variation in Ni/Mn enrichment in the least differentiated olivine (Figure

15). Mantle ranges are restricted to 2500-3000 ppm Ni and 500-1500 ppm Mn (Hoog et al., 29

2010; Foley et al., 2013) and differentiation trends toward lower Ni and higher Mn. NWRS

olivine have up to 3600 Ni and are consistently below 2000 ppm Mn, with a continuous range in

Ni/Mn ratios from > 4 to < 1. NVL mantle xenoliths have even higher Ni and lower Mn (~4000

and 150 ppm, respectively), encompassing a broader range of Ni/Mn ratios than NWRS olivine.

The broad range in Ni/Mn ratios for NWRS olivine and NVL mantle xenoliths suggests

heterogeneity in primitive melts. According to Foley et al. (2013), enrichment in Ni and Mn (i.e.

Ni/Mn < 1.5) is expected from melts that have incorporated recycled oceanic crust (i.e. OIBs),

whereas elevated Ni with little increase in Mn (i.e. Ni/Mn >2) may indicate melting in the

presence of metasomatic veins in continental lithosphere (Figure 15). This is relevant since

evidence for metasomatic influences (e.g., amphibole veins and phenocrysts in

peridotite/pyroxenite xenoliths) has been used to explain the genesis of NVL magmas (Coltorti et

al., 2004; 2007; Perinelli et al., 2006; 2011; Nardini et al., 2009; Gamble and Kyle, 1987;

Castillo et al., 2011).

The most primitive basalts from the NWRS fall on the spectrum from highly silica-

undersaturated basanite to moderately silica-undersaturated alkali basalt (Figure 2). The origin

of primitive alkaline magmas that have varying degrees of silica saturation has been explained by changes in degree of partial melting (Kay and Gast, 1973; Kogiso and Hirschmann, 2006;

Dasgupta et al., 2006), differences in source composition (Sun and Hanson, 1975; Marzoli et al.,

2000; Pilet et al., 2005; 2008) or early fractionation of more silica-rich phases (Green and

Ringwood, 1967). Basanite and alkali basalt in the NWRS form two sub-parallel trends when whole rock Mg# is plotted against Sr content (Figure 14b). This suggests that the magmas fractionated by similar processes, but with different starting compositions. Differentiation of basanite and alkali basalt along two separate magma lineages has often been invoked to explain 30

magma genesis within the WARS (LeMasurier and Thompson, 1990; Kyle et al., 1992; Wörner,

1999; Panter et al., 1997). Basanites erupted through continental lithosphere are slightly less

differentiated (Mg# 57-70; Ni = 106-338 ppm; Cr = 196-540 ppm) and less contaminated

(87Sr/86Sr < 0.7030; 187Os/188Os < 0.20) than alkali basalts (Mg# 56-68; Ni = 86-174 ppm; Cr =

179-477 ppm; 87Sr/86Sr > 0.7030; 187Os/188Os < 0.50) (Table 9); again similar to what has been

noted by Wörner (1999) and Panter et al. (1997).

For the least fractionated basalts, the degree of partial melting can be assessed using

ratios of highly incompatible trace elements to less incompatible elements (e.g., La/Yb and

Nb/Y), where lower degree melting will result in greater enrichment of highly incompatible

elements relative to moderate/low incompatible elements (e.g. higher Nb/Y and La/Yb). For

NWRS lavas, whole-rock ne-normative values and Sr contents show strong correlations to

degree of partial melting. Basanites are generally more enriched in highly incompatible elements

compared to moderate/ low incompatible elements (i.e. Nb/Y > 2.5), are more silica-

undersaturated (ne ≥ 9%), and have enriched Sr contents (> 800 ppm), which is consistent with

low degrees of partial melting of a garnet peridotite source. Alkali basalts correspond to

relatively higher degree melting (Nb/Y < 2.5), are less silica-undersaturated (ne ≤ 6 % to ~21%

hy) and have lower Sr (< 600 ppm) (Figure 16). Whole rock Zr contents plotted against indices

of differentiation (not shown) display the same relationship as shown for Sr in Figure 14b. The

difference in trace element concentrations over similar ranges in Mg# has been recognized by

Wörner (1999) for Cenozoic magmas in the WARS, and is attributed to different degrees of

partial melting. Previous studies of relatively unfractionated basalts from NVL have found

variations in radiogenic Sr-Nb-Pb isotope ratios that have been explained by heterogeneities in

the magma source (Wörner, 1999; Castillo et al., 2011). Similar compositional diversity in 31

ultramafic cumulates from NVL has been explained by varying contributions of metasomatic veins and adjacent depleted lithospheric mantle during the production of Cenozoic magmas

(Perinelli et al., 2011). The range of compositions measured in mantle xenoliths from different

areas in NVL are evidence of the heterogeneous nature of the upper mantle in this region

(Coltorti et al., 2004; Perinelli et al., 2006). The observed systematic increase in silica- undersaturation, Zr, Sr and La/Yb, Nb/Y, 143Nd/144Nd and 206Pb/204Pb ratios together with

decreasing 87Sr/86Sr from continental to oceanic domains is suggested by Castillo et al. (2011) to

reflect progressively smaller melt fractions from land to sea of a common heterogeneous mantle source. Partial melting of a heterogeneous lithospheric mantle has been proposed as the source of continental alkaline magmatism in the WARS and southwestern Pacific (Finn et al., 2005; Panter et al., 2006; Timm et al., 2010), and in other studies of intraplate alkaline magmatism (Lee et al.,

1996; Coltorti et al., 2007; Touron et al., 2008; Raffone et al., 2009) and is supported experimentally (Pilet et al., 2008; 2010; 2011).

18 An intriguing negative correlation exists between Nb/Y and δ Ool (Figure 17), which

appears stronger in alkali basalt (r2= 0.92) than basanite (r2= 0.54). The correlations suggest that

18 low δ Ool is related to low degrees of partial melting and therefore an indicator of source

18 composition. Low δ Ool in continental and oceanic basalts are often explained by assimilation

of hydrothermally altered mafic compositions from within the volcanic edifice or lithosphere

(Bindeman et al., 2008; Wang and Eiler, 2008). If low δ18O was inherited by contamination of

18 hydrothermally altered basalt, then δ Ool and Fo would be expected to be positively correlated

18 (Wang and Eiler, 2008). However, there is no correlation between Fo and δ Ool in NWRS

samples (Figure 8). Furthermore, crystallization temperatures for olivine in basanites and alkali

basalts are consistently > 1200°C, and by association with clinopyroxene, pressure greater than 32

0.9 GPa, indicating that olivine crystallized at depth (i.e. near the Moho) where contamination by

hydrothermally altered basalt is unlikely. It is well accepted that mass fractionation between

18O/16O at magmatic temperatures is insignificant, and therefore the degree of partial melting

18 cannot directly cause variation in δ Ool. Therefore in order to produce measurable differences in

δ18O by partial melting an initially heterogeneous δ18O source composition is required.

Oxygen isotope heterogeneity in the source may be related to metasomatism of the

lithospheric mantle during the amagmatic rift phase in the NWRS. Evidence for metasomatism at

the base of the continental lithosphere is provided by amphibole and amphibole-rich veins found

in mantle xenoliths from the WARS (Gamble and Kyle, 1987; Coltorti et al., 2004; Perinelli et

al., 2006; 2011). According to Perinelli et al. (2011), the metasomatic agent responsible for

isotopic disequilibrium between clinopyroxene and olivine was 18O-depleted relative to

ultramafic cumulates. Others have suggested that mantle metasomatism in the SW Pacific (New

Zealand and subantarctic islands) could be related to previous subduction (~500 to 100 Ma;

Panter et al., 2006), however subduction fluids/melts would likely result in 18O-enrichment of the

mantle (Eiler, 2001). Unfortunately, no current studies characterize the oxygen isotopic

composition of mantle derived fluids responsible for metasomatism.

5.2. Variations Across the Ocean-Continent Transect

This section focuses on variations in the least differentiated magmas from continent to

ocean and a discussion of T-P conditions of crystallization with implications for the architecture

of the lithosphere in the NWRS. The transect evaluates basanite and alkali basalt compositions

from the Adare Peninsula, Possession Islands and Adare Basin Seamounts (Figure 1) and

compares them with comparisons from previously studied NVL basalts, ultramafic cumulates, 33

and mantle xenoliths (Nardini et al, 2009; Armienti and Perinelli, 2010; and Perinelli et al., 2006;

2011).

Magmas begin to crystallize olivine and clinopyroxene at high pressure (0.7 to 1.3 GPa;

Figure 18), which is approximately 25 to 50 km depth assuming an average density for

continental crust of ~ 2.7g/cm3 (note that these pressure estimates are from clinopyroxene

collected within the CRS and OCTZ and are not from the oceanic domain). Previous workers

have estimated the depth of the Moho to be between 35 and 40 km beneath the Transantarctic

Mountains and between 18 and 20 km in the Ross Sea (Behrendt et al., 1991), therefore, the crystallization of olivine and clinopyroxene likely occurred at Moho depths or greater.

Presumably, mantle melts originated from within the subcontinental lithosphere (~130 to 160 km; Finn et al., 2005; Panter et al., 2006) and then rose to Moho depths where they stalled upon reaching the cooler crust where they began to crystallize. Kaersutite begins crystallizing at T-P

conditions at or above the Moho (1060°-1079°C and 0.6-0.8 GPa). The fact that kaersutite is

crystallizing at such high T-P conditions indicates that initial melts must have been volatile-rich,

becoming volatile-saturated shortly after crystallizing anhydrous phases (i.e. olivine and

clinopyroxene).

Temperatures and pressures estimated for olivine and clinopyroxene vary systematically

across the OCTZ. Olivine temperatures from the least differentiated magmas (i.e. alkali basalts

and basanites) increase from ~1200°C for CRS samples to ~1300°C for seamount samples in the

oceanic domain (Figure 19). Note that olivine temperatures derived from more differentiated

magmas (e.g., tephrite, hawaiite) are 50-100 °C cooler than alkali basalt and basanite. T-P

estimates from clinopyroxene exhibit an arching trend when plotted against distance from the

CRS (Figure 20), reaching minimum T-P conditions at ~20 km east of the CRS (Adare 34

Peninsula). The depression in T-P conditions of crystallization may correspond to a localized zone of lithospheric thinning in the NWRS which has been observed in seismic lines to the north and south of the NWRS (Trey et al., 1999; Lawrence et al., 2006). Lithospheric necking is supported by numerical modeling of extension in the WARS (Huerta and Harry, 2007). The necking zone is believed to have formed prior to the advent of volcanism, around the timing of focused extension in the Ross Sea (~43 to 26 Ma); it has been attributed to strengthening of the

West Antarctic lithosphere during syn-extensional thinning and cooling, ultimately leading to localized strain-weakening in the lithosphere (Huerta and Harry, 2007). Lithospheric necking in the NWRS is consistent with observed and modeled behavior of extension at magma-poor margins (Trey et al., 1999; Lawrence et al., 2006; Huerta and Harry, 2007). Likewise, the localization of volcanism along and adjacent to necking in NVL (i.e. Hallett Volcanic Province) is consistent with melting caused by shallow asthenospheric upwelling and subsequent melting at the base of the lithosphere. Furthermore, regional mantle flow models (Faccenna et al., 2008) provide a possible mechanism for the northward migration of late Cenozoic volcanism in this region, which decreases in age from the Melbourne Province (~15 Ma) to the Adare Basin (~3 to

<1 Ma; Granot et al., 2010; Panter and Castillo, unpublished).

Potential variations in mantle source compositions and partial melting processes across the continent-ocean transect are evaluated using whole rock compositions for the least fractionated basalts (Figure 21). Increases in silica-undersaturation and Nb/Y ratios indicate that compositions are controlled by lower degree melting with increasing distance ocean-ward from the CRS (Figure 21a and b). concentrations also increase with distance ocean-ward from the CRS (Figure 21c), suggesting control by partial melting but may also indicate difference in source composition. 35

Variations along the continent-ocean transect are also evaluated using mineral chemistryy and oxygen isotopes (Figure 21 d-f). The limited but distinct difference in Ni/Mn ratios for high

Fo olivine (≥ 85) from basanite and alkali basalt with distance across the transect supports a change in source composition from continent to ocean (<1.5 for basanite and > 2.5 for alkali basalt; Figure 21d). Recall that according to Foley et al. (2013), Ni/Mn ratios > 2 correspond to the continental-alkaline field and Ni/Mn < 1.5 corresponds to the OIB field (Figure 15). This difference is related to melting of metasomatized lithosphere in the continental case versus recycled oceanic lithosphere for OIBs.

With increasing distance from the CRS isotopes of oxygen for olivine and clinopyroxene also vary systematically (Figure 21e and f). The variation in oxygen isotopes with distance correlate to changes in degree of partial melting but must reflect source differences. As

18 discussed earlier, lower degree melts (i.e. high Nb/Y) correspond to lower δ Ool (Figure 17) and

Nb/Y increases with distance ocean-ward from the CRS (Figure 21a). If we assume that there is

a greater occurrence of metasomatic domains (i.e. amphibole/phlogopite bearing-veins) in the

continental lithosphere and that they are lighter in δ18O relative to the surrounding mantle, then it

seems reasonable that low degrees of melting will preferentially sample metasomatized mantle.

As degree of partial melting increases, the light δ18O signature becomes diluted as melting of the

higher δ18O adjacent lithosphere ensues.

5.3. Crustal Influences on Magmas

Textural and chemical evidence in many samples also indicate open system processes

(e.g., crustal contamination and magma mixing) as well as variation in magma ascent rates. This

is most likely influenced by the structural and compositional heterogeneity of the lithosphere 36

across the transect from continental to oceanic domains. Late Cenozoic volcanism in NVL and

the Adare Basin has been linked to steep normal faulting (Pliocene to Recent) which tend to align with early Miocene (~17 Ma) NE-SW trending structures (Granot et al., 2010). The decrease in thickness and nature of the lithosphere from MVP to HVP to Adare Basin would also exert some control on open system processes and magma ascent rates.

Contamination of magmas by continental crust is assessed using available whole rock

87Sr/86Sr, 143Nd/144Nd and 187Os/188Os isotope compositions (Hart and Kyle, unpublished;

Castillo et al., 2011; Juda et al., 2012). Contamination of magmas by continental crust is evident

by the high 87Sr/86Sr and 187Os/188Os ratios and low 143Nd/144Nd ratios in samples from the Malta

Plateau (Figure 22a). The Malta Plateau samples are hy- normative alkaline basalts (MA-009a

and MA-117) and slightly more differentiated (Mg# 56-59, Ni = 93-86 ppm, Cr = 272-349 ppm,

Sr < 600 ppm) relative to other samples (Figure 22b and c). The correlations between degree of

silica undersaturation and Sr content with 87Sr/86Sr and 143Nd/144Nd ratios (Figure 22b and c)

could be explained by assimilation of radiogenic continental crust with concurrent fractional

crystallization (AFC) but any further evaluation is unwarranted for only two samples.

Mixing between magmas may be recorded as complex compositional zoning in minerals

(i.e. reverse or oscillatory), which reflect changes in magma chemistry during crystal growth.

Resorption textures such as spongy clinopyroxene and sieve textures in plagioclase attest to

crystals that are not in equilibrium with the melt and may be antecrysts. The ubiquitous

occurrence of disequilibrium KD(Fe-Mg) values from olivine-liquid and clinopyroxene-liquid

pair calculations (Figure 10) is further evidence that foreign crystal debris (antecrysts or

xenocrysts) have been incorporated into magmas. This is most evident in basalts collected from the OCTZ, where spongy, zoned clinopyroxene is commonly found in concert with KD(Fe-Mg) 37

values that are greater than that expected for equilibrium conditions. Chemical transects across

zoned olivine further support magma mixing. For example, olivine from the Possession islands

exhibit an abrupt increased Fo and Ni contents (Figure 7b and e) which can best be explained by recharge and mixing of more mafic magma. Furthermore, most olivine exhibit a strong decreases in Fo at the rim (Figure 7a-c and e), which is unlikely to be caused by late-stage differentiation but is better explained if these represent antecrysts that have been entrained in a magma with a lower Mg/Fe ratio.

18 Many of the Δ Ocpx-ol fractionation values that indicate disequilibrium conditions (Figure

9a) may be the result of at least three possible senarios; 1) clinopyroxene crystallized in a lower

δ18O melt which mixed with a higher δ18O melt containing olivine, 2) olivine crystallized prior to the assimilation of a lower δ18O contaminant (e.g., hydrothermally altered basalt; Wang and

Eiler, 2008) followed by, or coincident with, clinopyroxene crystallization, or 3) olivine and

clinopyroxene crystallized in equilibrium, but are xenocrysts or antecrysts hosted in a lower δ18O melt in which xenocryst/antecryst clinopyroxene was isotopically reset (Mattey et al., 1994;

Perkins et al., 2006; Perinelli et al., 2011). The third scenario is proposed for NVL ultramafic

18 cumulates by Perinelli et al. (2011) to explain the diversity in Δ Ocpx-ol (+0.36 to -0.43 ‰). The

idea is that olivine retains a higher δ18O signature from earlier crystallization, while

clinopyroxene will equilibrate with the lower δ18O melt because the diffusion rate is twice as fast in clinopyroxene as in olivine (Farver, 1989). However, the overwhelming textural and compositional evidence for disequilibrium (presence of xenocrysts and antecrysts) supports that olivine and clinopyroxene did not crystallize in equilibrium with each other, favoring scenarios one and two. 38

Magmas that have ascended rapidly to the surface and were cooled quickly are characterized by quench textures such as skeletal olivine, variolitic crystal growth and glass-rich

(or crystal poor) groundmass. The limited amount of crystal fractionation, limited evidence for crustal contamination, the ubiquity of chemically homogeneous minerals as well as equilibrium mineral-melt compositions also support the rapid ascent of magmas with little or no interaction with the surrounding lithosphere. In the NWRS, these characteristics are dominantly observed in seamount lavas from the Adare Basin. Across the rift transect, timescales/magnitude for magmatic processes can be qualitatively assessed through diffusion of elements in olivine, where homogeneous grains correlate to short timescales or low magnitude change and intense zoning correlates to longer time scales or higher magnitude. Similarly, the reaction texture observed at amphibole rims has been attributed to instabilities caused by magma dehydration (Garcia and

Jacobsen, 1979) and is a function of magma ascent and potentially magma mixing (Rutherford and Hill, 1993) and can also be used in temporal assessments. The lack of amphibole in lavas at the CRS may be the result of 1) magmas with initially lower volatile concentrations which never reached saturation to crystallize amphibole, or 2) amphibole formed was destroyed by dehydration reactions on magma ascent to the surface.

A general model for magmatism across the continent-ocean transition is summarized in

Figure 23. Melting is initially produced at the base of the lithosphere from upwelling of the underlying asthenosphere triggered by extension within the Ross Sea. Differences in the degree of partial melting scavenge variable proportions of metasomatic veins and adjacent lithospheric mantle (Castillo et al., 2011). Mantle melts then rise to the Moho where they stall at the base of the cooler crust and begin crystallizing olivine and clinopyroxene. The depth of stalling varies across the transect and mimics the architecture at the base of the crust (i.e. changing thickness of 39 crust). Magmas erupted at the CRS travel through thick continental lithosphere, which lends to higher amounts of contamination and differentiation prior to eruption. Magmas erupted within the OCTZ travel through thinned continental lithosphere, which is perforated by a zone of lithospheric necking and intruded asthenosphere. Magmas crystallize over a range of temperatures and differentiated primarily by crystal fractionation and magma mixing. Oceanic magmas sampled at seamounts ascended rapidly to the surface and underwent the least amount of differentiation and contamination.

40

6. CONCLUDING STATEMENTS

Evaluation of mineral chemistry and oxygen isotopes from the least differentiated basalts in the NWRS expands upon previous studies which focus on whole-rock data and mantle xenoliths, providing new insights on the variation in mantle source compositions and partial melting processes across the continent-ocean transition. Magmas reflect varying levels of partial melting, differentiation and contamination evidenced by whole rock ne, Sr, Ni + Cr, Sr-Nd-Os isotopes, and supported by mineral chemistry which vary from continent to ocean.

1. Basalts in the NWRS have undergone limited differentiation and follow sub-parallel

differentiation trends suggesting fractionation occurred by similar processes with

different starting compositions (i.e. alkali basalt and basanite).

2. Variations in Ni/Mn ratios in high Fo (≥ 85 mol%) olivine may reflect heterogeneity in

the magma source.

18 3. Strong correlations between Nb/Y and δ Ool suggest that lower degree melts

preferentially consume a lower δ18O source interpreted as metasomatic veins in the

lithospheric mantle.

4. Crystallization of olivine and clinopyroxene began at or below the Moho and was

followed by amphibole in some samples at reasonably high T-P.

5. Lithospheric necking between the CRS and Adare Basin is supported by T-P estimates

from cpx-liq equilibrium crystallization.

6. Magmatic processes are exceedingly complex in the OCTZ and are likely related to

thinning and faulting within the extended continental lithosphere.

18 The correlation between δ Ool and degree of partial melting is perhaps the most intriguing finding in this study, suggesting that metasomatic veins are lower in δ18O than the adjacent 41

lithospheric mantle (since they would be more easily melted and dominate melt compositions at

low degrees of partial melting). However, the mechanism for producing a lower δ18O signature in the veins remains unkown. The composition and concentration of initial magmatic volatiles is poorly constrained for this region, lending to further ambiguity regarding the nature of metasomatic veining in the lithosphere and its influence on melting. The presence of kaersutite occurs in OCTZ and oceanic magmas and is a key to the T-P-X conditions of alkaline magmatism in the WARS, however investigations using this mineral are limited by current knowledge on the stability of calcic amphiboles (specifically in divergent settings). Until these issues are addressed, it may be difficult to say much more about the source of alkaline magmas in the WARS, however strong evidence for disequilibrium in OCTZ magmas (textural, compositional, and Δ18O between mineral phases) suggests there is much to be said about

physical controls on magmatism as it travels through the crust. Crustal influences on magmatism,

including an assessment of magma plumbing and timescales of magmatic processes may provide a unique view on rift architecture in not only the WARS, but magma-poor rift systems as a whole; an area of research that is yet to be exploited.

42

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Oppenheimer, C., R. Moretti, P. R. Kyle, A. Eschenbacher, J. B. Lowenstern, R. L. Hervig, and N. W. Dunbar (2011), Mantle to surface degassing of alkalic magmas at Erebus volcano, Antarctica, Earth Planet. Sci. Lett., 306(3), 261-271. Panter, K. S. (2007), Petrogenesis and source of lavas from seamounts in the Adare Basin, western Ross Sea: Implications for the origin of Cenozoic magmatism in Antarctica, paper presented at The 10th International Symposium on Antarctic Earth Sciences. Panter, K. S., S. R. Hart, P. Kyle, J. Blusztanjn, and T. Wilch (2000), Geochemistry of Late Cenozoic basalts from the : characterization of mantle sources in Marie Byrd Land, Antarctica, Chem. Geol., 165(3), 215-241. Panter, K., J. Blusztajn, S. R. Hart, P. Kyle, R. Esser, and W. McIntosh (2006), The origin of HIMU in the SW Pacific: evidence from intraplate volcanism in southern New Zealand and subantarctic islands, J. Petrol., 47(9), 1673-1704. Panter, K. and P. Castillo (2008), Petrology and source of lavas from seamounts in the Adare Basin, Western Ross Sea: Implications for the origin of Cenozoic magmatism in Antarctica, paper presented at AGU Fall Meeting Abstracts. Panter, K. S., P. R. Kyle, and J. L. Smellie (1997), Petrogenesis of a phonolite–trachyte succession at Mount Sidley, Marie Byrd Land, Antarctica, J. Petrol., 38(9), 1225-1253. Perinelli, C., P. Armienti, and L. Dallai (2011), Thermal Evolution of the Lithosphere in a Rift Environment as Inferred from the Geochemistry of Mantle Cumulates, Northern Victoria Land, Antarctica, J. Petrol., 52(4), 665-690. Perinelli, C., P. Armienti, and L. Dallai (2006), Geochemical and O-isotope constraints on the evolution of lithospheric mantle in the Ross Sea rift area (Antarctica), Contributions to Mineralogy and Petrology, 151(3), 245-266. Perkins, G. B., Z. D. Sharp, and J. Selverstone (2006), Oxygen isotope evidence for subduction and rift-related mantle metasomatism beneath the Colorado Plateau–Rio Grande rift transition, Contributions to Mineralogy and Petrology, 151(6), 633-650. Péron-Pinvidic, G. and G. Manatschal (2009), The final rifting evolution at deep magma-poor passive margins from Iberia-Newfoundland: a new point of view, Int. J. Earth Sci., 98(7), 1581-1597. Pilet, S., J. Hernandez, P. Sylvester, and M. Poujol (2005), The metasomatic alternative for ocean island basalt chemical heterogeneity, Earth Planet. Sci. Lett., 236(1), 148-166. Pilet, S., M. B. Baker, and E. M. Stolper (2008), Metasomatized lithosphere and the origin of alkaline lavas, Science, 320(5878), 916-919. Pilet, S., P. Ulmer, and S. Villiger (2010), Liquid line of descent of a basanitic liquid at 1.5 Gpa: constraints on the formation of metasomatic veins, Contributions to Mineralogy and Petrology, 159(5), 621-643. Putirka, K. D., M. Perfit, F. Ryerson, and M. G. Jackson (2007), Ambient and excess mantle temperatures, olivine thermometry, and active vs. passive upwelling, Chem. Geol., 241(3-4), 177-206. Putirka, K. D. (2008), Thermometers and barometers for volcanic systems, Reviews in Mineralogy and Geochemistry, 69(1), 61-120. Putirka, K. D., H. Mikaelian, F. Ryerson, and H. Shaw (2003), New clinopyroxene-liquid thermobarometers for mafic, evolved, and volatile-bearing compositions, with applications to lavas from Tibet and the Snake River Plain, Idaho, Am. Mineral., 88(10), 1542-1554. 47

Raffone, N., G. Chazot, C. Pin, R. Vannucci, and A. Zanetti (2009), Metasomatism in the lithospheric mantle beneath Middle Atlas (Morocco) and the origin of Fe-and Mg-rich wehrlites, J. Petrol., 50(2), 197-249. Rankenburg, K., J. Lassiter, and G. Brey (2005), The role of continental crust and lithospheric mantle in the genesis of Cameroon volcanic line lavas: constraints from isotopic variations in lavas and megacrysts from the Biu and Jos Plateaux, J. Petrol., 46(1), 169- 190. Ridolfi, F. and A. Renzulli (2011), Calcic amphiboles in calc-alkaline and alkaline magmas: thermobarometric and chemometric empirical equations valid up to 1,130° C and 2.2 GPa, Contributions to Mineralogy and Petrology, 1-19. Ridolfi, F., A. Renzulli, and M. Puerini (2010), Stability and chemical equilibrium of amphibole in calc-alkaline magmas: an overview, new thermobarometric formulations and application to subduction-related volcanoes, Contributions to Mineralogy and Petrology, 160(1), 45-66. Rocchi, S., P. Armienti, M. D’Orazio, S. Tonarini, J. R. Wijbrans, and G. Di Vincenzo (2002), Cenozoic magmatism in the western Ross Embayment: Role of mantle plume versus plate dynamics in the development of the West Antarctic Rift System, Journal of Geophysical Research, 107(10.1029). Rocholl, A., M. Stein, M. Molzahn, S. R. Hart, and G. Wörner (1995), Geochemical evolution of rift magmas by progressive tapping of a stratified mantle source beneath the Ross Sea Rift, Northern Victoria Land, Antarctica, Earth Planet. Sci. Lett., 131(3–4), 207-224, doi:10.1016/0012-821X(95)00024-7. Roeder, P. and R. Emslie (1970), Olivine-liquid equilibrium, Contributions to mineralogy and petrology, 29(4), 275-289. Rutherford, M. J. and P. M. Hill (1993), Magma ascent rates from amphibole breakdown: an experimental study applied to the 1980–1986 Mount St. Helens eruptions, Journal of Geophysical Research: Solid Earth (1978–2012), 98(B11), 19667-19685. Salvini, F., G. Brancolini, M. Busetti, F. Storti, F. Mazzarini, and F. Coren (1997), Cenozoic geodynamics of the Ross Sea region, Antarctica: crustal extension, intraplate strike-slip faulting, and tectonic inheritance, Journal of Geophysical Research, 102(B11), 24669- 24,696. Sawyer, D. S., M. F. Coffin, T. J. Reston, J. M. Stock, and J. R. Hopper (2007), COBBOOM: the continental breakup and birth of oceans mission, Scientific Drilling(5), 13-25. Sims, K. W. W., J. Blichert-Toft, P. R. Kyle, S. Pichat, P. J. Gauthier, J. Blusztajn, P. Kelly, L. Ball, and G. Layne (2008), A Sr, Nd, Hf, and Pb isotope perspective on the genesis and long-term evolution of alkaline magmas from Erebus volcano, Antarctica, J. Volcanol. Geotherm. Res., 177(3), 606-618. Sisson, T. and T. Grove (1993), Experimental investigations of the role of H2O in calc-alkaline differentiation and subduction zone magmatism, Contributions to Mineralogy and Petrology, 113(2), 143-166. Spear, F. S. (1981), An experimental study of hornblende stability and compositional variability in amphibolite, Am. J. Sci., 281(6), 697-734. Stern, T. A. and S. Uri (1989), Flexural uplift of the Transantarctic Mountains, Journal of Geophysical Research, 94(B8), 10315-10,330. 48

Storey, B. C., P. T. Leat, S. D. Weaver, R. J. Pankhurst, J. D. Bradshaw, and S. Kelley (1999), Mantle plumes and Antarctica-New Zealand rifting: evidence from mid-Cretaceous mafic dykes, Journal of the Geological Society, 156(4), 659-671. Stump, E. and P. G. Fitzgerald (1992), Episodic uplift of the Transantarctic Mountains, Geology, 20(2), 161-164. Sun, S. S. and G. N. Hanson (1975), Evolution of the mantle: geochemical evidence from alkali basalt, Geology, 3(6), 297-302. Timm, C., K. Hoernle, R. Werner, F. Hauff, P. v. den Bogaard, J. White, N. Mortimer, and D. Garbe-Schönberg (2010), Temporal and geochemical evolution of the Cenozoic intraplate volcanism of Zealandia, Earth-Sci. Rev., 98(1), 38-64. Touron, S., C. Renac, S. O'Reilly, J. Cottin, and W. Griffin (2008), Characterization of the metasomatic agent in mantle xenoliths from Deves, Massif Central (France) using coupled in situ trace-element and O, Sr and Nd isotopic compositions, Geological Society, London, Special Publications, 293(1), 177-196. Tréhu, A. M., A. Ballard, L. Dorman, J. Gettrust, K. Klitgord, and A. Schreiner (1989), Structure of the lower crust beneath the Carolina Trough, US Atlantic continental margin, Journal of Geophysical Research: Solid Earth (1978–2012), 94(B8), 10585-10600. Trey, H., A. K. Cooper, G. Pellis, B. della Vedova, G. Cochrane, G. Brancolini, and J. Makris (1999), Transect across the West Antarctic rift system in the Ross Sea, Antarctica, Tectonophysics, 301(1), 61-74. Valley, J. W. and N. T. Kita (2009), In situ oxygen isotope geochemistry by ion microprobe, MAC short course: secondary ion mass spectrometry in the earth sciences, 41, 19-63. Wang, Z. and J. M. Eiler (2008), Insights into the origin of low-< i> δ< sup> 18 O basaltic magmas in Hawaii revealed from in situ measurements of oxygen isotope compositions of olivines, Earth Planet. Sci. Lett., 269(3), 377-387. Weaver, B. L. (1991), The origin of ocean island basalt end-member compositions: trace element and isotopic constraints, Earth Planet. Sci. Lett., 104(2), 381-397. Weaver, S., B. Storey, R. Pankhurst, S. Mukasa, V. DiVenere, and J. Bradshaw (1994), Antarctica-New Zealand rifting and Marie Byrd Land lithospheric magmatism linked to ridge subduction and mantle plume activity, Geology, 22(9), 811-814. Wörner, G. (1999), Lithospheric dynamics and mantle sources of alkaline magmatism of the Cenozoic West Antarctic Rift System, Global Planet. Change, 23(1), 61-77. Zhang, H., D. Mattey, N. Grassineau, D. Lowry, M. Brownless, J. Gurney, and M. Menzies (2000), Recent fluid processes in the Kaapvaal Craton, South Africa: coupled oxygen isotope and trace element disequilibrium in polymict peridotites, Earth Planet. Sci. Lett., 176(1), 57-72.

49

FIGURES

Figure 1. Map of NWRS samples (ABS=Adare Basin Seamount, AP= Adare Peninsula, PI= Possession Islands, HP= Hallett Peninsula, DP= Daniell Peninsula, MP=Melbourne Local suite, MA= Malta Plateau, SI= Scott Island, and BI= Balleny Island). The continent-ocean transect (highlighted in yellow) includes samples from the AP, PI and ABS.

50

Figure 2. TAS diagram (after Le Bas et al., 1986) for selected NWRS volcanic rocks. Whole rock chemistry from Hamilton (1972), Johnson et al. (1982), Rocholl et al. (1995), Horning and Wörner (2003), Mortimer et al. (2007) and Panter and Castillo (unpublished). 51

Figure 3. Glomerocrysts and xenoliths in thin section: a) glomerocryst with euhedral to subhedral cpx and anhedral ol inclusions (xpl); b) glomerocryst with ol and zoned, kink banded cpx (xpl); c) anhedral cpx with interstitial glass (pl); d) ol + cpx ± plag (xpl). 52

Figure 4. Disequilibrium textures in thin section: a) deformed plag with sieve texture (xpl), b) plag with sieve texture (xpl), c) anhedral kearsutite with magnetite reaction rim (pl), d) clinopyroxene with spongy zone between core and rim (xpl).

53

Figure 5. Quench textures in thin section (found in ABS lavas from base of continental slope): A) variolitic cpx phenocrysts and skeletal ol, b) olivine phenocrysts and variolitic cpx microlites in glass-rich groundmass.

Figure 6. Mineral chemistry by zone: Mg-Ca-Fe quadrilateral showing olivine, clinopyroxene, and amphibole (shaded). Outliers: SCI-002 (Scott Island) is an evolved tephriphonolite, 24583 (Balleny Island), and MA-009a (Malta). Note that CRS cpx have the lower (≤ 50 wt%) Ca, oceanic cpx have higher (≥50wt%) Ca, and OCTZ cpx plot both above and below 50 wt% Ca. Olivine in the OCTZ has the greatest variatyion in FeO and MgO. No amphibole reported for CRS. 54

Figure 7. Olivine chemical transects: BSE images and corresponding chemical transects of six olivine chosen for detailed chemical analysis (CaO, MnO, NiO and Mg# = MgO/[MgO + FeO]*100). Transects begin in homogenous grain interiors (A= 0 µm) and continue across grain boundary with 1-5 µm step-size between measurements depending on nature of zoning: (a) MA- 009a d (CRS) with strong diffuse normal zoning at rim; (b) A227B d (OCTZ-Adare Peninsula) with oscillatory zoning; (c) PRR-5171 c (OCTZ-Possession Islands with strong, diffuse normal zoning at rim.

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Figure 7 (continued). (d) A232B d (OCTZ- Adare Peninsula) with weak, abrupt normal zoning at rim; (e) PRR-3872 a (OCTZ- Possession Islands) with diffuse reverse zoning and strong abrupt normal zoning at rim; (f) D12-1 c (OCTZ-ABS base of continental slope) with weak, abrupt normal zoning at rim. 56

18 Figure 8. δ Ool vs Fo. (a) All oxygen data (SIMS + LF) including averages and Tavg. Dashed lines and shaded bars represent dataset average and 2σ error of standards, respectively (green = SIMS, blue = LF). NVL ultramafic cumulates (Perinelli et al., 2011) and NVL basalts (Nardini et al., 2009) represented by dark gray and light gray regions, respectively. Note LF average is ~ 0.1 18 to 0.2 ‰ lighter than SIMS average. (b) SIMS δ Ool variations between core and rim. Note that rims are generally higher δ18O at Fo < 0.75. 57

18 Figure 9. Δ O fractionation for NWRS mineral pairs. Asterisks = NVL ultramafic cumulates, gray field = hydrous lherzolite (Chazot et al., 1997), black box is global average for mantle 18 peridotite (Mattey et al., 1994). (a) Δ Ocpx-ol; lines represent fractionation between cpx-ol at 18 1200 °C (+ 0.04 ‰), 800 °C (+ 0.08 ‰), and zero fractionation (Eiler, 2001). (b) Δ Oam-ol ; line indicates zero fractionation. Note that cpx is often lighter than coexisting olivine, where amphibole is more consistently close to olivine in δ18O and plots in similar ranges to that of Chazot et al. (1997). 58

Figure 10. Core-rim variations in calculated KD(Fe-Mg) exchange equilibrium: (a) olivine-liquid and (b) clinopyroxene-liquid by zone (filled circles = cores and open circles = rims). Whole rock chemistry used for liquid composition. Equilibrium ranges are ~ 0.30 ± 0.06 (2σ) for olivine (Roeder and Emslie, 1970) and ~ 0.27 ± 0.06 (2σ) for clinopyroxene (shaded areas; Putirka, 2008). A 2σ model error was used to account for the broad distribution in sample analyses and analytical error in measurements. Samples where both olivine and clinopyroxene are mostly above or below the equilibrium range, or bimodally distributed are highlighted in yellow. Samples where clinopyroxene is typically above equilibrium (though olivine is in equilibrium) are circled. 59

Figure 11. Mineral-liquid tests for equilibrium: A) Comparison of olivine Fo to whole rock Mg# (Rhodes diagram) which evaluates equilibrium as well as open/closed system behavior to explain disequilibrium (Rhodes et al., 1979). Solid and dashed line represent KD(Fe-Mg) equilibrium exchange between olivine and Whole rock composition (Roeder and Emslie, 1970). Arrows denote olivine removal, olivine accumulation differentiation and mixing . Note that disequilibrium olivine mostly follow differentiation and olivine removal trends. B) Observed vs predicted (from whole rock composition) cpx components for thermobarometry (Putirka 2008; 1999). Line represents 1:1 ratio between predicted and observed. Note that the DiHd shows the greatest deviation from the 1:1 line. Observed cpx tends to have higher DiHd than predicted, except for a few samples from the OCTZ. 60

Figure 12. Al-, Ti- Tschermak substitution in amphibole: Temperature sensitive Ti-Tschermak (A) and Pressure sensitive Al-Tschermak (B) substitution for NWRS samples compared to previously reported amphibole from NVL ultramafic cumulates and NVL, Melbourne Province (MP) and Foster Crater mantle xenoliths (Perinelli et al., 2011; 2006; Coltori et al., 2004; Gamble & Kyle, 1987, respectively). Note there is more deviation in the pressure sensitive model, but both show agreement between NWRS basalts and previously published data. Data regressed through NWRS data only (A and B). Regression in B excludes samples P74833 and SCI-002 which are outliers in the Pressure sensitive model (note these samples differ in composition where all other NWRS amphibole is kaersutite).

61

Figure 13. Amphibole P-T-H2O and thermal stability: NWRS basalts and NVL ultramafic cumulates (Perinelli et al., 2011) and xenoliths (Perinelli et al., 2006; Gamble and Kyle, 1987;

Coltori et al., 2004) based on Ridolfi et al. (2010); A) Temperature vs. pressure; B) H2Omelt vs temperature. Thermal stability minimum and maximum (dashed lines) are according to Ridolfi et al. (2010). Note strong agreement between most NWRS basalts and NVL ultramafic cumulates and xenoliths, both which plot outside suggested thermal stability. Samples P74833 and D4-3 have been labeled since they are outliers in B. 62

Figure 14. Differentiation trends: Alkali basalts and basanites represented by solid symbols, with differentiated lavas depicted by hollow symbols (arrows show differentiation). (a) Olivine and clinopyroxene fractionation denoted by decreasing Ni and Cr, respectively. (b) The increase in Sr with decreasing Mg# indicates that plagioclase is not fractionating from these magmas. Most primitive samples continental (P74794) and oceanic (D4-1) settings are labeled (which are also the most primitive alkali basalt and basanite, respectively).

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Figure 15. Ni/Mn ratios for high Fo olivine (≥ 85) after Foley et al. (2013). Fields denote global ranges for continental alkaline rocks and Mediterranean lamproites (CA) and ocean island basalts (OIB), with black box denoting narrow range in mantle values (Hoog et al., 2010; Foley et al., 2013). Arrow shows expected trend of differentiation (decreasing Ni/Mn). Dashed lines indicate ratio of Ni/Mn (i.e. 1:1, 2:1, and 4:1). Note that OIB are generally < 1.5 and CA > 2.0.NVL and Foster Crater mantle xenolith data from Perinelli et al. (2006) and Gamble and Kyle (1987), respectively. Note wide range in Ni/Mn ratios for NVL xenoliths.

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Figure 16. Partial melting trends: Relations between silica-undersaturation (ne) and whole-rock Sr with degree of partial melting (Nb/Y). Most primitive samples continental (P74794) and oceanic (D4-1) settings are labeled.

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Figure 17. Degree of partial melting (Nb/Y) vs. δ18O olivine. Note the strong negative correlation for alkali basalts and weaker negative correlation for basanites. Most primitive samples continental (P74794) and oceanic (D4-1) settings are labeled.

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Figure 18. T-P-Depth estimates for NWRS minerals (red= amphiblole, blue= clinopyroxene and green= olivine). Corresponding depths (right axis) are based on average density of continental 3 lithosphere (~2.7 g/cm ). Error bars for T-Pamp are within symbol size. Cpx T-P estimated by methods of Putirka et al. (2003), except for NVL mantle xenoliths (Armienti and Perinelli, 2010). Amp T-P estimated by methods of Ridolfi et al. (2010). Olivine T estimated from ol- liquid thermometer of Beattie (1993). Pressures from cpx-liq were used to plot olivine assuming the two phases crystallized at similar pressure. NVL ultramafic cumulates from Perinelli et al. (2011). Moho depth from Trehu et al. (1989), and East Antarctic craton geotherm after Armienti and Perinelli (2010). Note the paragenetic sequence, where olivine and clinopyroxene crystallize at/below the moho at near hydrous conditions, transitioning toward hydrous conditions prior to the crystallization of amphibole at/above the moho. 67

Figure 19. T-olivine with distance from the CRS: Olivine temperatures increase with distance from ocean-ward. Dashed line = continental rift shoulder and dotted line = base od continental slope. Temperatures from the most primative magmas (solid symbols) increase ocean-ward from the CRS. Lower temperatures correspond to more differentiated magmas (hollow sumbols). The two trends are offset by ~ 100 °C and may reflect multiple crystallization depths in lithosphere. 68

Figure 20. Clinopyroxene T-P with distance from the CRS. Dashed line = continental rift shoulder and dotted line = base of continental slope. Depths reported opposite of Pressure based on average density of continental lithosphere (2.7 g/cm3). Shaded area represents depth of Moho in coastal NVL (Trehu et al., 1989). Note the arching decrease in NWRS T-P between 0 and 40 km east of the CRS (curved lines).

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Figure 21. Geochemical variations with distance from the CRS: whole-rock (a-c) and olivine (d- f) across the continent-ocean transect. Dashed and dotted lines denote the CRS and base of the continental slope, respectively. A decrease in partial melting is shown by increasing Nb/Y (a). Increase in degree of silica-undersaturation shown by increasing ne (b); note the general composition of magmas shift from alkali basalt to basanite with increasing distance ocean-ward. Correlations between distance and Sr (c) may reflect changes in source or differentiation. Note that Ni/Mn ratios for olivine (d) are > 2.5 for alkali basalts near the CRS and < 1.5 for basanites 18 in the OCTZ. The negative correlation between δ Ool (e) and distance is as strong as Nb/Y (a) with distance, which suggests the two are dependent on each other (see Figure 16). There is a 18 weak negative correlation between δ Ocpx (analytical error is smaller than symbol size) and distance (f). 70

Figure 22. Assessment of contamination from continental lithosphere: evaluation across the continent-ocean transect using radiogenic Sr-Nd-Os isotopic ratios measured by Castillo et al. (2011) and Juda et al. (2012). Contaminated alkali basalts from Malta Plateau (circled) are further south than transect but plotted for contrast to OCTZ samples. Gray fields denote mantle values (e.g. contamination is negligible). Dashed and dotted lines represent CRS and base of continental slope, respectively. (a) Contamination with distance from the CRS. Note the agreement between all three systems which depict a decrease in contamination with distance ocean-ward from the CRS. Contamination and whole-rock ne-hy (b) and Sr (c). Note the wide range in ne-hy and Sr in relatively uncontaminated magmas. Sr and ne increase with contamination.

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Figure 23. Conceptual model for NWRS magmatism from continent to ocean. Dashed and dotted lines are approximated boundaries between CRS/OCTZ and OCTZ/oceanic domains, which are anticipated to be gradational. Depth of Moho as defined by Behrendt (1999) and Lawrence et al. (2006) and interpreted from continent to ocean by Huerta and Harry (2007). Depth of lithosphere/ asthenosphere boundary is approximated, but not well constrained for this region. Melting produced at the base of the lithosphere by decompression (bright red). Varying contributions of a heterogeneous source (i.e. metasomatic veining at the base of the lithosphere) are reflected in changes in degree of partial melting. Melts rise to the Moho where they stall at the base of the cooler crust and begin crystallizing olivine and clinopyroxene (dark red). The depth of crystallization varies across the transect and mimics the architecture (changes in thickness) of the lithosphere from continent to ocean.

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