Hafnium Isotope Geochemistry of the Gabbroic Crust Sampled Along the Mid-Atlantic

Home , Incompatible element, Isotope geochemistry, Lower oceanic crust

Ridge: Constraints on the Nature of the Upper Mantle

A thesis presented to

the faculty of

the College of Arts and Sciences of Ohio University

In partial fulfillment

of the requirements for the degree

Master of Science

Christine L. Thomas

August 2013

This thesis titled

Hafnium Isotope Geochemistry of the Gabbroic Crust Sampled Along the Mid-Atlantic

Ridge: Constraints on the Nature of the Upper Mantle

CHRISTINE L. THOMAS

has been approved for

the Department of Geological Sciences

and the College of Arts and Sciences by

Craig B. Grimes

Assistant Professor of Petrology

Robert Frank

Dean, College of Arts and Sciences

Abstract

THOMAS, CHRISTINE L., M.S., August 2013, Geological Sciences

Hafnium Isotope Geochemistry of the Gabbroic Crust Sampled Along the Mid-Atlantic

Ridge: Constraints on the Nature of the Upper Mantle

Director of Thesis: Craig B. Grimes

I characterize ɛHf values in zircon hosted in gabbroic rocks from three sites along

the Mid-Atlantic Ridge (ODP cores 1270D, 1275D and IODP core 1309D) in order to

examine mechanisms for igneous construction of the lower oceanic crust, and

characterize the isotopic heterogeneity of the underlying mantle source. The samples of

intrusive oceanic crust display a broad range of ɛHfzircon values from 13.3 to 22.1

(n=134), and zircons from individual drill cores spanning <1400 mbsf differ by up to 7.4

epsilon units. The variation within these single cores is similar to published data for

extrusive basalts sampled over scales of ~200 km. The distribution of ɛHf values within

cores supports crustal construction from individual, discreet pulses of magma. Since

isotopic variations cannot develop through igneous differentiation processes, the wide

range in values indicates derivation from a locally heterogeneous mantle source. Based on these observations, we propose that small-scale regions (<1 km) exhibit varying degrees of isotopic depletion, resulting either from differential melt extraction or refertilization in the geologic past. Considering geologically reasonable 176Lu/177Hf ratios

of natural magmas, the observed ɛHf range would requires the development of these compositionally distinct mantle parcels between 550 million to 2 billion years ago. 4

Dedication

This thesis is dedicated to my parents, David and Judy Thomas, for their support and understanding, and for reminding me to smile.

Acknowledgements

I would like to thank my adviser, Dr. Craig Grimes, for his help, guidance, and patience through my research and transfer to Ohio University. Without him this project would not have been possible. I would like to thank my committee members Dr. Doug

Green, and Dr. Greg Nadon for making me feel welcome at OU and always having his door open. I would like to acknowledge Dr. Joseph Wooden, Dr. Jorge Vazquez, Dr. Paul

Mueller, and Dr. George Kamenov for their help during analytical analyses and interpretation of results. Finally, I would like to thank the graduate students of Ohio

University for their support and encouragement.

Table of Contents

Page

Abstract……………………………………………………………………………………3

Dedication…………………………………………………………………………………4

Acknowledgments…………………………………………………………………………5

List of Tables……………………………………………………………………………...8

List of Figures……………………………………………………………………………..9

Chapter 1: Introduction…………………………………………………………………..11

1.1 Statement of Purpose………………………………………………………...11

1.2 Motivation for Research……………………………………………………..12

1.3 Overview...………………………….………………………………………..16

1.3.1 Mid-Ocean Ridge Processes……………………………………….16

1.3.2 Mantle Reservoirs………………………………………………….18

1.4 Background Geology………………………………………………………...20

1.4.1 15˚20' Fracture Zone, Mid-Atlantic Ridge………………………...21

1.4.2 ODP Hole 1270D, 14˚43' N………………………………………..22

1.4.3 ODP Hole 1275D, 15˚44' N………………………………………..23

1.4.4 Atlantis Massif, 30˚ N……………………………………………...24

1.4.5 Summary…………………………………………………………...27

Chapter 2: Methods………………………………………………………………………41

2.1 Zircon Significance and Separation Methods………………………………..41

2.2 The Lu-Hf Isotope System…………………………………………………...42

2.3 LA-ICP-MS Analytical Methods…………………………………………….43 7

2.3.1 Correction for isobaric interferences………………………………44

2.4 Zircon Trace Element Geochemistry via SHRIMP-RG……………………..47

Chapter 3: Results………………………………………………………………………..53

3.1 Zircon Textures in Cathodoluminescence………………….………………..53

3.2 Zircon Trace Element Geochemistry………………………………………...53

3.3 Zircon Lu-Hf Isotope Geochemistry………………………………………....55

Chapter 4: Discussions…………………………………………………………………...71

4.1 Intrasample Variation: Comparing Oceanic Zircon and the Hf-standard……71

4.2 Geographic Variations in Zircon and MORB ɛHf…………..……………….71

4.3 Growth of Lower Slow-Spreading Oceanic Crust…………………………...73

4.3.1 Heterogeneous Hafnium Isotopes along the MAR.………………..73

4.4 Investigating Mantle Source Compositions Using Zircon…………………...77

4.5 Variation in Crustal ɛHf by Mixing of Two End-Members ………………...79

4.6 Significance of ɛHf Variability………………………………………………83

4.7 Schematic Cross Section of the Mantle beneath the Mid-Atlantic Ridge…..85

4.8 Future Research……………………………………………………………...86

Chapter 5: Conclusions…………………………………………………………………..96

References………………………………………………………………………………..98

Appendix A: REE patterns for oceanic zircons with individual rock samples…………105

Appendix B: Histograms, probability density functions, error-weighted mean plots, and mean squared weighted deviates for all samples collected in this study……………….108

List of Tables

Page Table 1: SHRIMP-RG Trace Element Results for Zircons from 15˚20' Fracture Zone and

30˚N, Mid-Atlantic Ridge...……………………………………………………………...59

Table 2: LA-ICP-MS Isotope analysis of Zircons from 15˚20' Fracture Zone and 30˚N,

Mid-Atlantic Ridge………………………………………………………………….…...64

Table 3: ɛHf range and average for MAR sample sites (0-40% correction)………...... 69

List of Figures Page Figure 1: Model of proposed hypothesis showing incomplete mixing from a depleted

mantle reservoir and enriched/fertile small-scale reservoirs…………………………….28

Figure 2: Models for crustal accretion at ocean ridges…………………………………..29

Figure 3: Published Atlantic MORB ɛHf values along MAR latitudes…………….……30

Figure 4: Hf-isotope constraints on depleted MORB mantle (DMM)…………………...31

Figure 5: Topographic map noting MAR locations used in this study…………………..32

Figure 6: Seabeam bathymetric map of the region surrounding the 15˚20' Fracture

Zone……………………………………………………………………………………...33

Figure 7: Bathymetric profiles along cross sections for 1275D and 1270D………….….34

Figure 8: Representative rock types to host zircons……………………………………..35

Figure 9: Oblique view of Atlantis Massif………………………………………………36

Figure 10: Bathymetric profiles across Atlantis Massif…………………………………37

Figure 11: Plots of depth versus rock types, Mg#, and Pb/U ages………………………38

Figure 12: Rock types to host zircons for hole 1309D…………………………………..39

Figure 13: Large-format thin sections for the Southern Ridge of the Atlantis Massif…..40

Figure 14: Representative CL images of zircon grains from plutonic ocean crust……....49

Figure 15: Example screenshot of isotope spectra during a ~2.5 minute LA-ICP-MS analysis…………………………………………………………………………………...50

Figure 16: Percent isobaric interference correction versus ɛHf plots for Atlantis Massif

……………………………………………………………………………………………51 10

Figure 17: Percent isobaric interference correction versus ɛHf plots for 1270D and

1275D…………………………………………………………………………………….52

Figure 18: REE diagrams for zircons from sample sites………………………………...62

Figure 19: Trace element bivariate plots for U/Yb versus Y and Hf…………………….63

Figure 20: Comparison diagram of ɛ (0) and ɛ (t)………….……………………………68

Figure 21: Histogram of zircon ɛHf values collected in this study……………………....70

Figure 22: Published Atlantic MORB ɛHf and ɛNd values along MAR latitudes with the

averaged ɛHf and ɛNd values for 1270D, 1275D, 1309D, and Southern Wall………….87

Figure 23: Schematic drawing of crustal accretion models……………………………...88

Figure 24: Zircon ɛHf values collected in this study at the corresponding rock sample

depths and corresponding down-hole lithology……………………………………….…89

Figure 25: Averaged zircon ɛHf values and averaged zircon 206Pb/238U ages for each rock

sample in this study……………………………………………………………………... 90

Figure 26: Reference geochemical patterns for E-MORB and D-MORB……………….91

Figure 27: Zircon ɛHf values plotted against select trace element ratios………………..92

Figure 28: Modeling results depicting possible mixing between enriched and depleted mantle melts that had been previously fractionated ……………………………………..93

Figure 29: Modeling results depicting possible mixing between enriched and depleted mantle melts that had been previously fractionated.……………………………………..94

Figure 30: Idealized schematic illustrating the process of crust formation at sample sights

……………………………………………………………………………………………95

Chapter 1: Introduction

1.1 Statement of Purpose

The purpose of this study has been to characterize the hafnium isotopic

composition of multiply-intruded gabbroic oceanic crust sampled by ocean drilling near

the 15˚20’ Fracture Zone and Atlantis Fracture Zone (30˚N) to examine patterns of

vertical crustal growth and nature of the mantle source beneath the slow-spreading Mid-

Atlantic Ridge (MAR). Selected trace element ratios in zircon were examined to characterize geographic heterogeneity and investigate geochemical signatures reflective of mantle provenance. Published ɛHafnium isotope ratios are limited and exclusively measured from extrusive mid-ocean ridge basalts (MORB) recovered along the length of

the Mid-Atlantic Ridge. Basalts reportedly show large variations with latitude, and serve

as evidence for both isotopically enriched and depleted sources in the underlying mantle

(Andres et al., 2004). This study determines the first constraints from hafnium-isotope

ratios for single drill cores into the lower ocean crust, and provides constraints on the

local scale (1.4 km or less) variability within the crust and therefore the mantle source for

the igneous materials building the crust. Hafnium isotopes exhibit greater sensitivity than

conventional isotopic tracers like Sr and Nd, due to the shorter half-life of the 176Lu-176Hf

radiogenic system, and may provide new insights into the distribution of variably

depleted and enriched source material throughout the upper mantle in this region.

Previous work on a 1.4 km drill core at the 30˚N site has shown multiple igneous intrusive contacts indicating average intrusion thicknesses of ~10 m (John et al., 2009), 12

variable Pb/U ages in zircon, and chemical differentiation trends indicative of a

protracted history of igneous activity. Similar geologic relationships are reported for

gabbroic rocks recovered at site 1275D. At site 1270D, thin (cm-scale) dikes intrude mantle peridotite. Thus, the crust appears to have been built by thin pulses of melt, rather than by accretion from a large magma chamber in which melts become homogenized before cooling to form the crust. These sites contrast with modern fast-spreading ridges, in which the crust is thought to grow from a long-lived, narrow melt-lens (i.e., the gabbro glacier; Phipps-Morgan and Chen, 1992); small, local scale heterogeneities in the mantle source would likely be homogenized during magma mixing and replenishment of a long- lived melt lens. I hypothesize that the small-scale intrusive melts building the gabbroic crust at the 15˚20’ Fracture Zone and Atlantis Fracture Zone will preserve heterogeneous hafnium isotopes reflecting initial heterogeneities in the underlying mantle (Figure 1).

1.2 Motivation for Research

The Mid-Atlantic Ridge (MAR) is a slow-spreading oceanic ridge system

(spreading rates ~25mm/yr) along which extension is accommodated by magmatism and faulting. Sampling of the seafloor in several locations, including those targeted for this study, reveals a complex lithospheric architecture with lower crustal gabbros and mantle peridotite exposed directly on the seafloor. These rock types have been uplifted from estimated depths of up to 7 km along large-offset normal faults (Tucholke and Lin, 1994;

Canales et al., 2007; Grimes et al., 2011; Schoolmeesters et al., 2012). These deep exposures offer new opportunities to examine the development of the intrusive igneous 13

crust that covers nearly two-thirds of the Earth’s surface as well as the mantle from which

it originates. Extrusive basalts have traditionally been the means by which the mantle

source has been characterized globally. Based largely on studies of basalts erupted at both

mid-ocean ridges and ocean island settings and melting experiments, it is generally

agreed that the upper mantle is comprised of harzburgite (depleted peridotite) and the

lower mantle is comprised of incompatible-element enriched lherzolite (Zindler et al.,

1984; Wilson, 1989). In addition to modeling of the residual source from basalt

compositions, rare exposures of upper mantle on the seafloor and in ophiolites confirm

depleted harzburgite as the dominant rock type in the upper mantle (Kelemen et a., 2004;

Cannat 1996).

The general, simplest, model for the Earth’s mantle is a two-layer model in which

the depleted upper mantle (MORB source) and enriched lower mantle (OIB source)

convect independently (Basaltic Volcanism Study Project, 1981); however in detail the

mantle is undoubtedly more complex. Recent isotopic and trace element studies of both

igneous and mantle material suggest that the upper mantle is not uniformly mixed

(Harvey et al., 2006; Portner et al., 2011). One limitation to characterizing the upper

mantle from erupted basalt is that in magmatically robust environments (like the fast-

spreading East Pacific Rise), melts from a large region of the mantle may coalesce in a

shallow crustal magma chamber and become homogenized. Additionally, in areas where

high degrees of partial melting occur (i.e., fast-spreading ridges), small-volume components within the mantle might get diluted. Compared to East Pacific Rise basalts

(which have a more uniform and less radiogenic Hf isotopic composition), the MAR 14

basalts record more variable and more radiogenic Hf isotopic compositions (Salters et al.,

2011). Along the MAR, systematic isotopic variations have been recognized from basalts

sampled along the scale of an entire ridge (~10,000’s kms) and correlations between

multiple isotope systems are noted on the scale of individual ridge segments (~100’s

kms); the heterogeneity within the mantle constrained thus far is interpreted to represent

geochemical variations on similar scales (100s to 1000s km) (Dosso et al., 1993; Andres

et al., 2004). Although systematic changes in the mantle are known, questions remain

about the regional and local magnitude of isotopic heterogeneity in the upper mantle, and

growing evidence suggests the mantle is heterogeneous on smaller scales (<1-10’s km)

(Harvey et al., 2006; Portner et al., 2011). Furthermore, pervious isotopic research on abyssal peridotites within the 15˚20’ FZ region reflects varying degrees of melt extraction and refertilization through time and supports the existence of local-scale heterogeneity

(Harvey et al., 2006).

Recent drilling near the 15˚20’ Fracture Zone (FZ) and Atlantis Massif (30˚N) where the volcanic crust is thinned or absent reveals discontinuous gabbroic crust intruding mantle peridotite. The lower crust and mantle has been exposed on the seafloor by large-offset normal faults. Geochemical studies of peridotites recovered near the

15˚20’ FZ indicate that they represent some of the most depleted harzburgites (olivine +

orthopyroxene rock) recovered along the MAR (Harvey et al., 2006). Peridotites from the

Ocean Drilling Program Hole 1274A were found to exhibit variable and highly non-

radiogenic osmium isotope ratios (187Os/188Os = 0.114-0.126), indicating that these

abyssal peridotites went through melt depletion event in the past (up to 2 Ga) and do not 15

relate to the average degree of depletion in the upper mantle that is sampled by typical

MORB which represents an aggregate of sources dominated by more chondritic

compositions (Harvey et al., 2006). Depleted peridotites with highly non-radiogenic

osmium-isotope ratios contrast with limited (n = 3) ɛHf isotope analyses of basalts from

the ridge segments north and south of the 15˚20’ FZ, which give normal MORB to

enriched MORB-like trace element and isotope signatures (Dosso et al., 1991 & 1993;

Agranier et al., 2005). The peridotite and basalt examined in the region so far therefore

cannot be complimentary residue-melt pairs. An isotopically more enriched mantle

source must be present at depth. Collectively, these observations indicate local scale

heterogeneities within the column of melting mantle beneath the 15˚20’ FZ.

Multiply-intruded gabbroic sections sampled by drilling and dredge provide a way

to evaluate the mantle source from which melts were derived. Unlike extrusive basalts,

which may homogenize during ascent or due to residence in the shallow melt lens, core

observations from the three sites in this study reveal cm- to m-scale intrusive bodies within peridotite or larger gabbro plutons (John et al., 2009). These small-scale intrusive bodies, emplaced over 100s kyr or more, may preserve a more robust measure of the isotopic composition of the direct mantle source. Zircons hosted in these gabbroic samples are ideal for investigating primary chemical information of the magmas; zircons are highly resistant to modification by hydrothermal alteration and they concentrate a variety of elements that serve as isotopic tracer systems, especially hafnium, allowing relatively precise measurement on single crystals.

1.3 Overview

1.3.1 Mid-Ocean Ridge Processes

Mid-ocean ridge magmatism is one of the most obvious expressions of plate tectonics. Horizontal extension brings hot, uprising mantle melts from depth to the colder seafloor surface; it cools, accretes to the base of the spreading plates, and forms new oceanic crust. Partial melting of the mantle at spreading centers is the mechanism for the creation of oceanic crust. As the mantle convects upward beneath mid-ocean ridges adiabatically, the pressure decreases and eventually mantle peridotite intersects its solidus

(the transition from complete solid to partial melt) and melts progressively during further ascent. The more buoyant magmas rise to the surface and create a basic sequence of crustal rock types observed in the ocean crust.

The textbook view (Penrose model; Anonymous, 1972) of the igneous ocean crust consists of a carapace of extrusive basaltic lava flows and pillow lavas erupted directly onto the seafloor, underlain by shallow intrusive sheeted dikes, intrusive gabbro and related cumulate rocks that crystallized directly from a steady-state magma chamber, and ultramafic rocks of the lower crust and mantle (Klein, 2003) (Figure 2a). While drilling in the East Pacific Rise has shown that this structure is typical of oceanic crust formed at fast spreading ridges (Wilson et al., 2006), a substantial percentage of the crust generated at slower spreading ridges does not follow this architectural pattern.

Slow-spreading mid-ocean ridges are defined as having a full spreading rate of

<55 mm/yr and represents approximately 60% of the global ridge system by length

(Schwartz et al., 2005; Dick et al., 2006). The crust at slow-spreading ridges is thinner 17

(Chen, 1992; Cannat, 1996; White et al., 2001), and in several locations the crustal

architecture does not fit the standard Penrose model (Dick et al., 2006). Decreased

spreading rates have been correlated with increased tectonic (less magmatic) extension

that results in roughening of basement topography, decreased crustal thickness, and the

presence of gabbro and peridotites exposed directly on the seafloor (Lizarralde et al.,

2004). Although few ridge segments have been thoroughly mapped, up to 50% of the

segment south of the 15˚20’ FZ is estimated to comprise gabbro and peridotite

(Schroeder et al., 2007). Slow spreading ridges like the MAR also exhibit a greater

degree of isotopic heterogeneity than fast spreading ridges; this is attributed to more

effective magma mixing within an axial melt lens within fast-spreading ridges and less

effective magma mixing in the absence of an axial melt lens within slow-spreading ridges

(Batiza, 1984).

Large tracts along some segments of slow spreading ridges, such as the 15˚N area of the MAR, developed without abundant volcanism and are described as avolcanic

(Schroeder et al., 2007); there the seafloor is composed of gabbroic crust and mantle peridotite brought to the surface by faulting instead of extruded basalt (Lagabrielle et al.,

1998) (Figure 2b). For slow-spreading systems, spreading rate may have a greater effect on melt extraction than on total melt production (Lizarralde et al., 2004), but this does not necessarily require lower degrees of melt production. Rather, the melt may be ineffectively extracted from the mantle, due to the presence of a thick, cool, brittle lithospheric lid. Thus, melt gets distributed over a wider vertical interval, intruding as smaller accumulations of melt (Cannat, 1996). This results in more isolated pockets of 18

melts, which do not accumulate and homogenize in a high level magma chamber at the

base of the crust (Blackman et al., 2006; Delacour et al., 2008).

1.3.2 Mantle Reservoirs

The presence of multiple discrete mantle reservoirs is evident mostly from trace

element (REE) compositions of basalts as well as long-lived radiogenic isotope

systematics within mantle-derived melts, and recently, mantle peridotite itself. These

geochemical parameters have led to the distinction of both enriched and depleted end-

member mantle reservoirs, where ocean island basalts (OIB) and normal mid-ocean ridge basalts (N-MORB) generally define the enriched and depleted end members (Stracke et al., 2003). Even though enriched and depleted end-members are often defined, it is also possible that the mantle spans a range of intermediate compositions.

Development of a Depleted Mantle (DM) reservoir begins by extraction of incompatible elements from the upper mantle by partial melting; those melts then ascend to form ancient oceanic crust, which is subsequently subducted back into the mantle. The result is a chemically-zoned upper mantle showing strong depletion of highly incompatible elements (Saunders et al., 1988). Following silicate melt extraction, a depleted peridotitic residue is left (Saunders et al., 1988). Partial melting leads to well- known fractionation of several parent-daughter isotope pairs, including Rb-Sr, Sm-Nd,

Lu-Hf, and U-Pb. The parent isotopes Sm and Lu behave more compatibly (concentrated in the residue left after partial melt is removed) relative to their respective daughters, whereas Rb and U behave relatively incompatibly (concentrated in the partial melt). The fractionation of parent and daughter isotopes has resulted in the Earth’s crust and mantle 19

having different parent-daughter ratios, which over time, results in very different isotopic

ratios of the daughter elements. As a result of ancient melt extraction, DM has therefore

developed characteristically high 143Nd/144Nd, 176Hf/177Hf and low 87Sr/86Sr, 206Pb/204Pb

(Zindler and Hart, 1986; Saunders et al., 1988; Workman and Hart, 2005) relative to

enriched mantle or crustal reservoirs.

The Enriched Mantle (EM) reservoir refers to either fertile mantle (mantle that

has not previously undergone partial melting) or mantle that has been refertilized by the

addition of new melt (Andres et al., 2004). EM is characteristically enriched in incompatible elements and exhibits low 143Nd/144Nd, 176Hf/177Hf, and high 87Sr/86Sr,

207Pb/206Pb, 208Pb/206Pb (Zindler and Hart, 1986; Saunders et al., 1988; Andres et al.,

2004).

This project uses these contrasting isotopic systems as tracers to investigate the

isotopic variation recorded by multiply-intruded gabbroic crustal sections. As expected,

measured values fall within two defined end-members based on MORB along the MAR

(Andres et al., 2004): the E-MORB reservoir (MORB generated from EM) and the N-

MORB reservoir (MORB generated from DM). In crustal rocks, the primary reservoir for

the element hafnium is the mineral zircon (typically 1-2 wt % hafnium); such high

concentrations can be precisely measured on single crystals. The 176Hf/177Hf ratios in

zircons determined in this study are compared to established values for the global average

depleted MORB mantle as well as its enriched and depleted end-members (Workman and

Hart, 2005), and also to previously published MAR MORB ɛHf values (Figures 3, 4). 20

On the scale of the MAR, ɛHf values vary by a greater magnitude than that of

more conventional isotopic tracers like Nd because Lutetium (176Lu = 37.1 Gyr) has a shorter half-life than samarium (147Sm = 106 Gyr) (Kinney and Mass, 2003). Chemically,

Sm and Nd are very similar (both are rare earth elements) and partition to a lesser degree,

while Lu (a rare earth element) and Hf (a high field strength element) are more different

and partition to a higher degree. Thus, Hf isotopic values may provide greater resolution

of isotopic domains (Chauvel and Blichert-Toft, 2001). Fractionation of Lu-Hf during partial melting of the mantle has led to development of different isotopic ratios in depleted versus enriched mantle sources on the earth today. ɛHf along with trace element data are used to distinguish mantle reservoirs that have been separated for long periods of

time.

1.4 Background Geology

Drilling near the 15˚20’ Fracture Zone and Atlantis Massif (30˚N) (Figure 5)

reveal highly fractured crust, comprised of variably evolved gabbro (troctolite, gabbro,

olivine gabbro, gabbronorite, Fe-Ti oxide gabbro) and mantle peridotite. These sites were

selected for drilling because the volcanic crust (i.e., basalt) is thinned or missing

altogether, allowing deep exposures to be sampled. Numerous igneous contacts observed

in the cores, along with geochronology and geochemical evidence, suggest that the

sampled crustal sections were built by intrusion of multiple discrete pulses of magma

over 100’s kyr (Blackman et al., 2006; Grimes et al., 2008; Grimes et al., 2011).

1.4.1 15˚20’ Fracture Zone, Mid-Atlantic Ridge

The 15˚20’ Fracture Zone has been the subject of several studies since it was recognized in the 1990’s that deep crustal and mantle exposures crop out in this area

(Dosso et al., 1993; Cannat et al., 1996; Kelemen et al., 2004; Harvey et al., 2006).

Exposures of peridotite and gabbro are found at both the north and south of the fracture zone. In part to understand the distribution of these rock types, Leg 209 of the ODP drilled 19 holes at 8 drill sites in the area (Kelemen et al., 2007). This is a good place to study isotopic characteristics because, unlike fast spreading ridges like the East Pacific

Rise, basaltic crust is thin to absent.

Leg 209 of the Ocean Drilling Program (ODP) recovered drill cores along the

Mid-Atlantic Ridge, from 14˚43’N to 15˚44’N (Figure 6a), where residual mantle peridotite intruded by gabbroic rocks resides on the seafloor (Kelemen et al., 2007). This magma-starved area is in the vicinity of the 15˚20’ Fracture Zone (FZ) that offsets by

~175 km one of the slowest portions of the ridge (mean full spreading rate of 25km/m.y.)

(Fujiwara et al., 2003; Godard et al., 2008).

A geochemical gradient is observed in MORB sampled along the northern MAR

(Dosso et al., 1993). The basalt compositions evolve from E-MORB in the southern 14˚N region to N-MORB in the northern 16˚N region (Dosso et al., 1991, 1993).

Based on whole rock, trace element, and Re-Os isotope geochemical analysis, peridotites from ODP Hole 1274 are extremely depleted; they underwent a prolonged period of partial melting and melt extraction over the past ~1.5 Gyr (Harvey et al., 2006; Godard et al., 2008). 22

1.4.2 ODP Hole 1270D, 14˚43’N

Core 1270D (total depth = 57 mbsf) is located on a westward-dipping (~10˚) slope ~15 km east of the MAR axial valleys (Figure 6c). It is part of a laterally continuous detachment fault scarp that did not result in the formation of a domal complex

(Fujiwara et al., 2003; Schroeder et al., 2007).

About 91% of the core is composed of harzburgite, with minor amounts of dunite

(7%) and gabbroic dikes (2%) (Kelemen et al., 2004) (Figure 7b). The peridotites are intruded by several generations of thin dikes, which have subsequently been altered under

low grade (< 400˚C) metamorphism (Kelemen et al., 2004). The dikes are 1-2 cm thick and preserve evidence of highly localized, synkinematic, high-temperature alteration

(Bach et al., 2004; Jöns et al., 2009). Within the dikes, relict igneous minerals are limited to zircon, apatite, amphibole, and Fe-Ti oxide minerals (Jöns et al., 2009). Geochemical modeling, along with the presence of zircon and apatite, concludes that these dikes are likely altered trondhjemite intrusions rather than truly ‘gabbroic’ in composition (Jöns et al., 2009).

Two of these zircon-bearing dikes intruding harzburgite were sampled at 19 and

25 mbsf (less than 6m apart) (Figure 8c, 8d). Samples 1270D-19 and 1270D-25 comprise cm-scale greenschist-grade dikes that are dominantly chlorite-actinolite or tremolite schist (Jöns et al., 2009). Zircon 206Pb/238U ages yielded 1.28±0.03 Ma for 1270D-19 and

1.14±0.04 Ma for 1270D-25; the ages reflect magma intrusion spanning at least >70 kyr

(Grimes et al., 2011).

1.4.3 ODP Hole 1275D, 15° 44’ N

Core 1275D (total depth ~209 mbsf) is located ~28 km west of the MAR axis and

drilled on a domal, bathymetric high interpreted as an ocean core complex (MacLeod et

al., 2002; Fujiwara et al., 2003; Kelemen et al., 2004) (Figure 6b). The surface of the dome has been interpreted as the footwall to a large-offset normal fault with >15 km displacement (MacLeod et al., 2002; Fujiwara et al., 2003; Kelemen et al., 2004).

Seventy four percent of Core 1275D is composed of gabbroic rock types

(including gabbronorite, Fe-Ti oxide gabbro, and minor olivine gabbro), 14% ultramafic rocks, 10% diabase, and 2% plagiogranite (Kelemen et al., 2004) (Figure 7a). The upper

~50m of core consists primarily of ultramafic rock types with minor diabase while the lower ~160m is composed of gabbroic rocks intruded by diabase, plagiogranite, and other gabbroic units (Kelemen et al., 2004). Numerous igneous contacts within the core imply that the gabbroic section was built by multiple intrusive events (Kelemen et al., 2004).

Five cm-scale leucocratic dikes were sampled from core 1275D, spanning 10-200

mbsf (Figure 8a, 8b). Four samples (at 134, 144, 180, 200 mbsf) came from the

predominantly gabbroic lower section of the core; one sample (at 10 mbsf) was taken

from the predominantly ultramafic upper section of the core. The small-volume dikes occur throughout the core and are interpreted to represent the latest melts to crystallize; they typically comprise plagioclase ± amphibole ± quartz ± Fe-Ti oxide ± alkali feldspar

± zircon ± apatite (Kelemen et al., 2004).

Previous zircon 206Pb/238U ages for 1275D yield constrains on the duration of accretion. Ages indicate that the numerous intrusive pulses of magma observed in the 24

core occurred over a period of 40 kyr (Grimes et al., 2011). The five rock samples used in

this study give ages that are indistinguishable within analytical uncertainties; they range from 2.12 ± 0.17 to 2.31 ± 0.10 Ma (Grimes et al., 2011).

1.4.4 Atlantis Massif, 30˚ N

Samples at this location were collected from both IODP drill core U1309D and the southern wall of Atlantis Massif (Figure 9). Cores for site U1309 were drilled during

Expedition 304/305 of the Integrated Ocean Drilling Program. Site U1309 is located on the central dome of the Atlantis Massif (a well-studied oceanic core complex) 14–15 km west of the MAR axis; it is interpreted to be a gently sloping, corrugated detachment fault surface that has exhumed lower crustal and mantle rocks (Blackman et al. 2006). Core

U1309D was drilled into the displaced footwall of the detachment fault and penetrated

1415.5 m below the seafloor (Blackman et al., 2006).

Eighty four percent of the core is composed of predominantly gabbroic rock types, 7% of Fe-Ti oxide gabbro, ~5% ultramafic rocks, 3% diabase, and <1% felsic dikes (Blackman et al., 2006) (Figure 10, 11a). The core has been grouped into three lithostratigraphic units (Figure 11b). Unit 1, extending from the seafloor to 600 mbsf, shows relatively uniform whole-rock Mg # (MgO/FeO+MgO). Unit 2, extending from

~600-1235 mbsf, shows a systematic decrease in Mg # with height. This trend is

consistent with a typical magmatic differentiation series, involving fractionation of Mg-

rich minerals like olivine resulting in Fe-enrichment in later-formed melts. Unit 3, extending from ~1235 mbsf to the bottom of the hole (1415 mbsf), shows variable whole- 25

rock Mg # and reflects more fractionated rock types (Fe-Ti oxide gabbro and felsic dikes) that are found in the base of the overlying unit (Grimes et al., 2008).

Seven rocks were sampled in hole 1309D between at 58, 250, 564, 820, 1175,

1327, and 1415 mbsf (Figure 12), and spanning all three lithostratigraphic units. D58 comprises an altered leucocratic (amphibole-plagioclase) zircon-bearing vein intruding gabbro. D564, D820, D1327 are enriched in Fe-Ti oxides and occur as randomly distributed patches in undeformed, coarse-grained gabbro. Samples D250 and D1175 is also a Fe-Ti oxide enriched, but reside near undeformed, sharp contacts with gabbro

(Blackman et al., 2006). D1415 consists of cm-scale dikes that are made up of anorthosite, quartz diorite, and tonalite leucocratic melt intrusions (Grimes et al., 2008).

206Pb/238U isotopic dating reveals variation between the three lithologic groups mentioned above that represents real temporal variation; this indicates that emplacement of the entire section was not done in one short-lived period of accretion but in multiple intrusive events (Figure 11c) (Grimes et al., 2008).

Existing 86Sr/87Sr ratios for whole rocks within hole 1309D show a range from

0.70261 to 0.70904; elevated Sr isotope compositions are restricted to ultramafic lithologies from the top of the core are interpreted to arise from alteration by seawater

(Delacour et al., 2008). Whole rock analysis of the gabbroic rocks define a narrow range in Sr isotope (0.70261-0.70370) and Nd isotope (9.1-11.5) compositions, which are consistent with an origin from a depleted mantle source with homogeneous Sr and Nd isotope compositions (Delacour et al., 2008). This observation will be further evaluated 26

in this study with new hafnium isotope ratios from multiple zircons within each whole

rock.

The southern ridge of the Atlantis Massif (Figure 9) forms the wall of the Atlantis

Transform and rises ~700 m below sea level (Blackman et al., 2006). Mapping by submersible suggests that approximately 70% of this area is peridotite; the remaining

30% comprises gabbro as small cm- to m- scale intrusive bodies (Blackman et al., 2002).

Three surficial samples (3652-1333, 3646-1205, 3647-1359) and one dredge (D3-21) were analyzed as part of the present study (Figure 13). These samples comprise highly altered, variably deformed mafic-ultramafic rocks that were likely intruded by evolved, late-stage melts (Grimes et al., 2008). Zircon is typically found in altered veins/dikes, but also occurs in the schistose matrix of highly deformed samples as in 3646-1205 (Grimes et al., 2008). Lead-uranium zircon ages for the southern ridge samples become older further from the ridge axis, as expected from simple seafloor spreading (Grimes et al.,

2008).

Strontium (87Sr/86Sr) isotope ratios measured on the southern ridge have slightly

higher values than seen in 1309D and indicate a higher degree of seawater alteration

(Delacour et al., 2008). Reported ɛNd values also indicate modification by seawater-rock

interaction, and range from mantle-like to less radiogenic (Delacour et al., 2008). Thus, the existing datasets demonstrate that primary igneous isotope ratios of whole rocks have been obscured. It is anticipated that zircon will preserve igneous isotopic compositions more reliably than previous analyses of whole rock.

1.4.5 Summary

Samples from all three sites studied consist of small-volume gabbro intrusions,

occurring either as discrete dikes into peridotite or within larger complexes of gabbroic

crust. The crust in these settings did not form from a large, steady state melt lens in which

melts delivered from the underlying mantle mix and become homogenized. While Nd-

isotopes from hole 1309D suggest an origin for melts from a homogeneous, depleted

mantle reservoir, isotope ratios from whole rocks along the southern wall are altered by

hydrothermal processes. Determining ɛHf in zircons hosted in these intrusive melts will

further constrain the isotopic character of the crust-forming magmas, and allow an

investigation of the nature of their source in the mantle. The style of crustal accretion in

these areas is most consistent with a many-sill model (Kelemen et al., 2004; John et al.,

2009; Grimes et al., 2011), indicating the gabbroic crust grew by injection of small sills

over a protracted period of time. The 176Hf/177Hf isotope data provide constraints on the

homogeneity the source of these melts. If the source is a deeper, larger magma chamber,

or an isotopically homogeneous mantle, values will be similar. Alternatively, if the

source is more reflective of a many-sill model, isotope values will show greater variation compared to an isotopically homogeneous mantle.

Figure 1: Model of proposed hypothesis showing incomplete mixing from a depleted mantle reservoir (light green) and enriched/fertile small-scale reservoirs (dark green/yellow). Based after Jeffcoat (2012).

Figure 2: Models for crustal architecture at ocean ridges after accretion. A. Classic interpretation of the Penrose Model for a fast spreading ridge based on the Oman Ophiolite. B. Model for the anomalous 14˚N-16˚N area of the MAR. Modified from Dick et al. (2006).

Mohns- Knipovitch Ridge

Figure 3: Published Atlantic MORB ɛHf values along MAR latitudes. For reference, ɛHf values are shown for E-DMM (11.1; 2σ enriched from the average), Average DMM (16.8), and D-DMM (25.3; 2σ depleted from the average), respectively. ɛHf values were calculated from 176Hf/177Hf values in Workman and Hart, 2005 using BSE 176Hf/177Hf of 0.282725 [Bouvier et al., 2008].

Published ɛHf values were calculated from 176Hf/177Hf values in Chauvel and Blichert- Toft (2001); Andres et al. (2004); Agranier et al. (2005); Debaille et al. (2006); and Pet DB using BSE 176Hf/177Hf of 0.282725 [Bouvier et al. (2008)]. Sample localities for this study near the 15˚20’ FZ are between 14˚33’N and 15˚42’N. Localities near the Atlantis Massif are between 30˚N to 30˚17’N. Arrows indicate data near the Iceland hotspot (60˚- 70˚N) and the Mohns-Knipovitch Ridge (~72˚N), which is not representative of typical MORB-source magmas.

Figure 4: A. histogram showing the ɛHf values for the depleted MORB mantle (DMM). ɛHf values for E-DMM, Average DMM, and D-DMM are the same as in Figure 3. ɛHf values were calculated from 176Hf/177Hf values in Workman and Hart, 2005 using BSE 176Hf/177Hf of 0.282725 [Bouvier et al., 2008]. B. Histogram of previously published ɛHf values measured on basalts in the vicinity of sample locations used in this study. These are limited to only 3 analyses for the 15˚20’region, and 4 analyses from the 30˚N area.

Published ɛHf values were calculated from 176Hf/177Hf values in Chauvel and Blichert- Toft (2001); Andres et al. (2004); Agranier et al. (2005); Debaille et al. (2006) using BSE 176Hf/177Hf of 0.282725 [Bouvier et al., 2008]. MAR localities range from ~44˚S to ~80˚N, with areas affected by the Iceland plume (60˚N-70˚N) omitted due to their high degrees of enrichment. 32

Figure 5: Map of the Atlantic Ocean indicating the location of samples used in this study. The locations of the Atlantis Massif and 15˚20’ Fracture Zone are noted with yellow stars.

Figure 6: A. Seabeam bathymetric map of the region surrounding the 15˚20’ Fracture Zone (modified from Fujuwara et al., 2003) B. Blowup bathymetry at ODP site 1275; note spreading direction parallel corregations on the surface of smooth, domal bathymetric high or core complex (modified from Fujuwara et al., 2003) C. Blowup bathymetry at ODP site 1270; the ridge observed is a single large fault surface and not a core complex (modified from Fujuwara et al., 2003) 34

Figure 7: Bathymetric profiles along cross section lines A-A’ and B-B’ (See figure 6). Vertical exaggeration on both cross sections is 2:1. Downhole columns depict lithologies recovered at various depths in A) ODP core 1275D and B) ODP core 1270D; recovered rock types are averaged over 10 m intervals and normalized to 100%. Stars indicate the depth of samples analyzed in this study (yellow stars indicate localities of hand samples in figure 8). Modified from Grimes et al. (2011).

Figure 8: Representative rock types that host zircons. Scale bars are all 2 cm. A) and B) are felsic dikes seen in hole 1275D that typically comprise plagioclase ± amphibole ± quartz ± Fe-Ti oxide ± alkali feldspar ± zircon ± apatite; samples pictured are located at depths of 144 and 180 mbsf. C) and D) are centimeter-scale greenschist-grade dikes (dominantly chlorite-actinolite or tremolite schist) that intrude peridotite seen in core 1270D; samples picture are located at depths of 19 and 25 mbsf. From Grimes et al., (2011).

Figure 9: Oblique view of Atlantis Massif (looking WNW). IODP Hole U1309D is indicated on the corrugated detachment fault surface. Surficial samples (3652-1333, 3646-1205, 3647-1359, D3-21) used in this study are located on the southern wall. The base map was created using Generic Mapping Tools (GMT) software (Wessel and Smith, 1998). Figure is from Grimes et al. (2008).

Figure 10: Bathymetric profile for cross section A-A’ (see figure 9) across Atlantis Massif, with no vertical exaggeration. Lithologic column depicts the running average of rock types recovered in hole 1309D. White on the right-hand side of the lithologic column indicates no recovery. Stars indicate locations of samples (seen in figure 12) used in this study. Modified from Grimes et al. (2008); Blackman et al. (2006).

Figure 11: Plots of depth versus A) 20m running average of rock type recovered in Hole U1309D. White indicates no recovery. B) Whole rock Mg#, along with inferred lithologic supergroups and fault zones. C) 207Pb and 230Th corrected 206Pb/238U sample weighted average zircon ages. Shaded fields represent Jaramillo and Cobb Mountain normal polarity intervals [Cande and Kent, 1995]. Figures modified from: Blackman et al. (2006); Grimes et al. (2008).

Figure 12: Rock types host to zircon for hole 1309D. Scale bars are 4 cm. A. is an altered leucocratic (amphibole-plagioclase) vein intruding gabbro; B., C., D., E., and F. are Fe-Ti oxide gabbros. G. is a cm-scale plagiogranite dike made up of anorthosite, quartz diorite, and tonalite leucocratic melt intrusions.

Figure 13: Large-format thin sections of surficial samples from the Southern Ridge of Atlantis Massif. Scale is 1.5 in. A. D3-21: metasomatized amphibolite+chlorite+talc-rich fault rock with relict olivine and orthopyroxene. B. 3646-1205: talc-amphibole schist. C. 3647-1359: brecciated amphibole-chlorite rock with talc pseudomorphing orthopyroxene. D. 3652-1333: foliated tremolite+chlorite-rich rock with coarse relict Fe-Ti oxide. From Grimes et al. (2008)

Chapter 2: Methods

2.1 Zircon Significance and Separation Methods

Zircon (ZrSiO4) is a common accessory mineral found throughout the geologic

record; it is most abundant within intermediate and felsic igneous rocks but is also found

in mafic igneous, metamorphic, and sedimentary rocks. The high resistance of zircon to

chemical and physical modification during geologic events (metamorphism, magmatism,

sedimentation) makes it ideal for investigating primary magmatic age and chemical

information. In addition, single crystals reflect the isotopic composition of the parent magma, whereas several lines of geochemical and petrological evidence indicate that

cumulate gabbroic rocks may form from multiple igneous events or experience melt-rock reactions during magmatic accretion below the ridge axis (Lissenberg and Dick, 2006).

To accomplish the proposed goal of characterizing the hafnium isotope composition of multiply-intruded gabbroic crust, individual zircons have been analyzed in situ using ion microprobe techniques. SEM-Cathodoluminescence (CL) imaging was conducted prior to analysis to characterize chemical zoning, and identify cracks, inclusions, or altered domains that were then avoided during analysis (Figure 14).

Zircons from the 15˚20’ Fracture Zone and Atlantis Massif were previously separated from whole rocks using standard crushing and mineral separation techniques and then mounted in epoxy and polished (Grimes et al. 2007). SEM-CL imaging using the Zeiss EVO 50 scanning electron microscope was completed at Mississippi State

University’s Electron Microprobe Center in Spring 2012. 42

2.2 The Lu-Hf Isotope System

Radioactive lutetium (176Lu) decays by beta emission to radiogenic hafnium

(176Hf) with a half-life of 37.1 Gyr (λ176 = 1.867e-11 yr-1) (Kinny and Maas, 2003). The

long half-life means that the system does not provide precise age determination on the

young samples as studied here, but instead acts as a geochemical tracer for identifying

distinct mantle sources that have been separated over very long time scales. Lutetium is

more compatible in the mantle than Hf, and thus when melting occurs, Hf is more highly

fractionated into the melt. The resulting mantle residue has a higher Lu/Hf ratio, which

leads to enrichment of radiogenic 176Hf over time relative to crustal reservoirs, or bulk

silicate earth. Hafnium isotope ratios are often reported using epsilon notation (εHf),

which provides a direct, whole number comparison to the ratio within the bulk earth

(BSE, or CHUR). The term εHf is defined as:

176 177 0 176 177 0 4 εHf = {[( Hf/ Hf) sample / ( Hf/ Hf) CHUR ] – 1} * 10 (Eqn #1)

where 177Hf is the nonradiogenic isotope of Hf and “CHUR” (chondritic uniform

reservoir) denotes chondritic values 176Hf/177Hf. Positive values of εHf reflect

enrichment of radiogenic 176Hf relative to CHUR and are indicative of residence in a

source with a higher Lu/Hf ratio than CHUR (i.e., the depleted mantle). Negative values

of εHf indicate depletion of radiogenic 176Hf.

The most compatible element in zircon, besides viiiZr4+ (IR= 0.84 Å), is viiiHf4+ (IR

= 0.83 Å). Typically zircon will possess 0.5-2.0 weight percent Hf (Hoskin and

Schaltegger, 2003), and the Hf concentrations collected from zircons in this study range 43

from 0.57 to 2.1 weight percent (Table 1). However, viiiLu3+ (IR = 0.977 Å) is less

compatible in zircon and is not present in high concentrations (Hoskin and Schaltegger,

2003). The difference in compatibility results in parent-daughter fractionation of Lu and

Hf by zircon. Heavy isotopes do not fractionate from one another during magmatic processes, and therefore the Hf isotope ratios recorded by zircon will represent the Hf isotope ratios of the melts from which the zircon crystallizes. Furthermore, the low Lu/Hf ratio that is characteristic of zircon (this study’s averaged Lu/Hf value is 0.016) results in

negligible ingrowth of 176Hf relative to total 176Hf since crystallization.

2.3 LA-ICP-MS Analytical Methods

Hf isotope concentrations were collected in June 2012 using the Nu Plasma multi-

collector LA-MC-ICP-MS at The University of Florida and used the protocol outlined by

Mueller et al. (2007). Isotope analyses were performed with on-line Lu and Yb isobaric

interference corrections, using 176Lu/175Lu = 0.02653 and 176Yb/172Yb = 0.5870, both

within the range of published values (Mueller et al., 2008; Vervoort et al., 2004). All

isotope ratios were corrected for mass bias using 176Hf/177Hf = 1.46718 (Mueller et al.,

2008). Analyses of FC-1 zircon standard were used to monitor the precision of the

instrument and yielded an average of -22.6 ɛHf (176Hf/177Hf = 0.28215 ± 0.00005, 2 SD; n

= 60) with an external error of 1.4 epsilon units (2 SD during the analytical session. The

FC-1 standard data are within error of multi-grain dissolution analyses of FC-1

(176Hf/177Hf = 0.282174 ± 0.000013, 2 SD) (Mueller et al., 2008). The 176Lu decay

constant of λ176Lu = 1.867x10E-11 (Söderlund et al., 2004) was used in all calculations. 44

LA-MC-ICP-MS collects Hf isotopes by continuously scanning the sample across

a range of masses; this creates a range of spectra over a ~2.5 minute time interval (Figure

15). Hf isotopic measurements were made on Faraday detectors acquiring 180Hf, 178Hf,

177Hf, 176Hf +176Yb + 176Lu, 175Lu, 174Hf + 174Yb, 173Yb, 172Yb, and 171Yb simultaneously.

Because different isotopes of equal mass (i.e., isobars) are present in the zircons, analyses

have to be corrected for isobaric interferences.

2.3.1 Correction for Isobaric Interferences

Typical isobaric interferences for continental and synthetic zircon are 20% or less

(Kamenov et al., 2010). However, one of the characteristic features of zircon from

oceanic crust is high concentrations of HREE. For example, the average Yb/Hf ratio for

zircons hosted in porphyry copper deposits in northern Chile is 0.04 (±0.04 1SD; n=210;

Ballard et al., 2002) whereas the average Yb/Hf ratio collected from oceanic zircons in

this study is 0.13 (±0.14, 1SD, n= 133). Zircons in this study have a higher-than-normal

correction that need to be assessed.

Previous studies have shown that in situ Hf isotopic measurements by LA-ICP-

MS typically require small 176Lu and 176Yb isobaric interference corrections. A study of

synthetic, doped zircons showed that accuracy and precision decrease systematically with increasing REE/Hf ratios (Kamenov et al., 2010). Kamenov et al. (2010) showed that isotope ratios requiring less than 20% correction for isobaric interferences could be made without decreasing the accuracy of ɛHf. Due to the characteristically high HREE content of oceanic zircons, larger corrections in isobaric interferences are common. Because of 45

the high concentration of elements with 176Hf isobars, an overall evaluation of the data

quality was first performed to evaluate potential analytical bias.

Measured 176Hf/177Hf ratios were first examined by comparing them to the percent

external correction required for 176Lu and 176Yb isobaric interference (Mueller et al.,

2008). The percent correction is determined using the following calculation:

176Hf/177Hf (raw)-176Hf/177Hf (corrected) /176Hf/177Hf (raw) (Eqn #2)

Corrected 176Hf/177Hf ratios were determined by deducting Lu and Yb isobaric interferences (using 176Yb/172Yb = 0.5870 and 176Lu/175Lu = 0.02653; Vervoort et al.,

2004) and corrected for mass bias using 178Hf/177Hf = 1.46718.

In order to assess the effects of 176Yb isobaric interferences, percent corrections were compared to the calculated ɛHf values for each analysis (Figures 16, 17).

Considering all samples taken from the Atlantis Massif, the isobaric interference correction data (for Hf-corrected 176Hf/177Hf ratios) define a ɛHf range of 13 – 25 epsilon

units from 0% to 40% correction. As the percent correction increases above 40%, ɛHf

data scatter to both higher and lower values and the ɛHf range increases from -1.6 to 49

epsilon units (Figure 16a). Based on this observation, data points with >40% correction

are considered less accurate with the result that 28 isotope analyses from the Atlantis

Massif were omitted (Figure 16b). The low R2 values for both scatter plots in Figures 16a

and 16b (0.0008 and 0.0546, respectively) indicate that there is real scatter in the data,

and this could potentially be related to geochemical heterogeneity.

When plotted against the percent isobaric interference correction, ɛHf values from

ODP core 1270D show that the percent correction have significant scatter with three data 46

points separated by an increase of ~10 % correction (Figure 17a). While these data points do not change the ɛHf range, the two samples over 40% correction (1270D-19-9 [41%] and 1270D-19-15 [46%]) were omitted to retain consistency between data sets. For all

ɛHf values from ODP 1275D, isobaric interferences are under 40% correction and none were omitted from the data set (Figure 17b). Again, the low R2 values for 1270D and

1275D (0.191 and 0.0599, respectively) indicate that there is real scatter within the data

set. Based on this observation, data points with >40% correction are considered less

accurate and 2 isotope analyses from ODP core 1270D are omitted. Thus, a total of 30

(out of 174) isotope analyses were omitted because of high (>40%) correction.

As a secondary screen of data quality, 176Hf/177Hf internal errors were evaluated

for each analysis. “Internal error” refers to the variance or reproducibility in 176Hf/177Hf

during a single spot analysis (which typically lasts 2.5 minutes). Nine isotope analyses (1

from 1270D, 4 from 1275D, 4 from 1309D) were discarded because of high 176Hf/177Hf

internal error of ± 0.00013-0.00024 (1 SD) compared to the average internal error of ±

0.000024 (1 SD) for these data. The discarded internal errors on measured 176Hf/177Hf are

approximately an order of magnitude greater (10-4) than the rest of the observed values

(10-5-10-6). Large variance is an indication of accidental analysis of inclusions or

intersection of isotopically-distinct domains in a zircon during the laser ablation process.

Beam intensity is another possible cause of variations in internal error. Since these

analyses are considerably less precise than remaining data, they were omitted.

There was also one zircon that displayed analytical inconsistencies not associated

with percent isobaric interference or 176Hf/177Hf internal errors. During analysis, laser 47

burned through zircon 1270D-19-09a (Table 2) very quickly, which resulted in a

shortened analysis time (less than 2.5 minutes) and greater analytical variability.

Although this grain has low internal error and low percent isobaric interference, it is

discarded because of its anomalous analysis.

2.4 Zircon Trace Element Geochemistry via SHRIMP-RG

In addition to hafnium, zircons contain many other trace elements record

information about the parental magmas, and possibly the initial source of those magmas.

Trace element analyses for zircons from the 15˚20’ FZ and Atlantis Massif were

performed in July 2011 using the sensitive high-resolution ion microprobe (SHRIMP-

RG) located at the U.S. Geological Survey- Stanford Ion Probe Laboratory and used the

protocol outlined in Grimes et al. (2009). In SHRIMP analyses, a primary oxygen beam

loosens ions from the zircon grain and the material is then extracted into the mass

spectrometer and analyzed using a single detector. Zircon spot analyses were performed

using a 3-6 nA primary beam current and 15-20 µm spot size while working at ~11,000

mass resolution (M/ΔM). Analyses consisted of a single cycle that peak step sequentially

from 9Be+ to 254UO+. The concentration standard CZ3 (a Sri Lankan megacryst, Ireland and Williams 2003) was run after every ten unknowns.

To evaluate whether geochemical differences exist in zircons from different geographic locations, comparison of trace element ratios in zircon from the slow-

spreading 15˚20’ Fracture Zone and Atlantis Massif (Figures 5, 6, 9) were conducted.

Geographic differences might reflect variations in mantle source composition. These data 48

also add to the small, but growing database of zircon trace element geochemistry useful for provenance studies in other geologic settings.

Figure 14: Representative cathodoluminescence images of zircon grains from plutonic ocean crust. Scales are 100µm. Areas targeted for LA-ICP-MS analysis are noted by white circles. Small craters are visible within several of these circles; the craters are from previous trace element analysis preformed using SHRIMP-RG. Oceanic zircons typically exhibit subhedral to euhedral, tabular to equant crystal habit. Zoning in CL is limited (light/weak) to absent and roughness around grain edges is due to polishing. Black spots within the grains are mineral inclusions. Grains are largely homogeneous in CL contrast and are interpreted to have formed from a single-stage growth event.

Figure 15: An example screenshot of isotope spectra during a ~2.5 minute LA-MC-ICP- MS analysis. During analysis, the zircon is ablated and the isotopic composition of the ablated material is continuously monitored. The resultant ablated material is drawn into a vacuum, and ionized using inductively coupled plasma (ICP) with a temperature of 5000- 10000 degrees. The ions are then drawn into the mass spectrometer, separated by massing using electrostatic lenses and a magnetic sector, and then measured using a multi- collector allowing simultaneous analysis of the elements shown in this screen shot. The intensity of the ion signal is shown by the y-axis. Intensity always decreases over time because as the ablation pit in the zircon gets deeper, less material escapes into the vacuum. Each color spectra represents a different mass. The isotopes 180Hf, 178Hf, 177Hf, 176Hf +176Yb + 176Lu, 175Lu, 174Hf + 174Yb, 173Yb, 172Yb, and 171Yb were acquired simultaneously.

51 A. Atlantis Massif: All Data 70

40 f

H 30 ɛ 176/177 Hf corr. 20 R² = 0.0008 Linear (176/177 Hf corr.) 10

-10 0% 20% 40% 60% 80% 100% % Correction

B. Atlantis Massif: Data < 40% Correction 70

40 f H ɛ 176/177 Hf corr. 30 Linear (176/177 Hf corr.) 20 R² = 0.0546 10

0 0% 20% 40% 60% 80% 100% % Correction

Figure 16: Percent correction versus ɛHf for all zircons analyzed from Atlantis Massif, 30˚N MAR. Considering all the data, corrections varied from ~5% up to a maximum to ~85%, with the higher values reflecting zircon grains with high corrections of Yb (up to 5132 ppm). For all samples taken from Atlantis Massif, percent isobaric interference correction vs ɛHf plots show that the Hf and Yb correction data cluster within the same ɛHf range (13-25) below about 40% correction, but become scattered at corrections >40%. This is interpreted to result from large deviations in accuracy for larger correction factors. The trend line displays a R2 value of 0.0008. A. As the percent correction increases from 40%, data start trending downward. B. When data with a correction greater than 40% is removed, the downward trending is also removed. The trendline displays a R2 value of 0.0546. 52 A.

ODP 1270D 30

20 f

H 15 ɛ 176/177 Hf corr. 10 R² = 0.191 Linear (176/177 Hf corr.)

0 0% 10% 20% 30% 40% 50% % Correction

ODP 1275D 25 R² = 0.036 20 R² = 0.0599 15 f 176/177 Hf corr. (10 mbsf) H ɛ 176/177 Hf corr. (130-200 mbsf) 10 Linear (176/177 Hf corr. (10 mbsf)) Linear (176/177 Hf corr. (130-200 mbsf)) 5

0 0% 10% 20% 30% 40% 50% % Correction

Figure 17: Percent correction versus ɛHf for all zircons analyzed from ODP 127D and ODP 1275D. A. Data for ODP core 1270D. Hf and Yb correction data vary from 8% to 46% correction, with all but three data points having less than 24% correction. A linear regression has a R2 value of 0.2. While the two data points over 40% correction do not change the ɛHf range for this rock sample, they were omitted to retain consistency between data sets. B. Data for ODP core 1275D. Hf and Yb isobaric interference correction data are under 40% isobaric interference correction and none were omitted from the data set. Both linear regressions for 10 mbsf and 130-200 mbsf display a R2 value of 0.04 and 0.06, respectively. 53

Chapter 3: Results

3.1 Zircon Textures in Cathodoluminescence

Scanning Electron Microscopy – Cathodoluminescence (SEM-CL) imaging

reveals areas in each zircon free of mineral inclusions or cracks prior to analysis, which

were then targeted for trace element and isotope analysis. CL imaging of zircon samples

show that zoning within grains is limited (light/weak) to absent. Grains exhibit subhedral

to euhedral, and tabular to equant crystal habit; roughness around the edges is due to

polishing. Black spots within the grains are mineral inclusions (Figure 14).

3.2 Zircon Trace Element Geochemistry

New zircon trace element (TE) data were collected for five samples from OPD core 1275D, and compiled along with previously published TE data from the 15˚20’ FZ and Atlantis Massif (Grimes et al, 2009). Both sets of TE data are reported in Table 1.

Chondrite- normalized rare earth element (REE) patterns for oceanic zircons recovered

from the 15˚20’ FZ and Atlantis Massif are shown in Figure 18. The patterns display

trends typical of igneous zircon, with enrichment from the light rare earth elements

(LREE) toward the heavy rare earth elements (HREE), and Eu- and Ce- anomalies that occur in all grains analyzed here (Figure 18; zircon REE patterns for individual rocks are given in appendix A). These patterns are characteristic of igneous zircon (Hoskin and

Ireland, 2000). Substitution of these elements is largely controlled by simple or coupled

substitution mechanisms involving Zr and Si. The 0.84-Å radius of the [8]Zr4+ (8-fold 54

coordination and 4+ ionic charge) is more closely matched by smaller-radii HREE (e.g.,

[8]Lu3+, 0.977 Å) than larger-radii LREE (e.g., [8]La3+ , 1.160 Å), so LREE are generally

incompatible in the zircon structure (Hoskin and Schaltegger, 2003). Eight-fold

coordinated Ce4+ has an ionic radius of 0.97 Å that is close to Lu (0.977 Å); Ce is

therefore more compatible to the zircon structure than other LREE and creates a positive

anomaly spike in the REE diagrams (Hoskin and Schaltegger, 2003). Eight-fold coordinated Eu2+ has a large ionic radius of 1.25 Å relative to [8]Zr4+ (0.84 Å); Eu is

therefore more incompatible and creates a negative anomaly (Hoskin and Schaltegger,

2003). Studies have shown that zircon REE patterns from a wide range of crustal rock

types and tectonic settings are very similar; making the patterns not useful indicators of

provenance (Grimes et al., 2007).

Trace elements used in graphical analysis include U (5 – 7556 ppm), Yb (210 –

7065 ppm), Hf (5760 – 21487 ppm), Y (350 – 20497 ppm), Nb (1 – 175 ppm), and P

(169 – 5263 ppm) (Table 1).When U/Yb is plotted against Hf and Y, zircons from the

15˚20’ FZ form a tight cluster between 500-4000 ppm Y, 7000-10000 ppm Hf, and 0.03-

0.3 U/Yb; variation is on the order of one magnitude of less (Figures 19a-b). The 1270D

average values are Hf = 8548 ppm (± 1680 1SD), Y= 1280 ppm (± 366 1SD), and U/Yb

= 0.1 (± 0.09 1SD). The 1275D average values are Hf = 9040 ppm (± 740 1SD), Y =

1634 ppm (± 820 1SD) and U/Yb = 0.08 (± 0.02 1SD). More variability is seen within the

Atlantis Massif data, which scatter from 200-20,000 ppm Y, 6000-20,000 ppm Hf, and

0.01-1.0 U/Yb; variation is on the scale of two orders of magnitude in U/Yb and Y, and

no greater than one order of magnitude in Hf. (Figures 19a-b). The Atlantis Massif 55

average values are Hf = 11744 ppm (± 2893 1SD), Y = 7270 ppm (± 7237 1SD), U/Yb =

0.13 (± 0.17 1SD).

Based on the average values and standard deviations mentioned above and on

visual observation of Figures 19a-b, it is clear that the Atlantis Massif zircon

compositions are distinct from the 15˚20’ FZ. Zircons analyzed from Atlantis Massif

show a general positive correlation between U/Yb ratio with Hf and Y; this is interpreted

to result from fractional crystallization of the host magmas.

3.3 Zircon Lu-Hf Isotope Geochemistry

Hafnium isotope concentrations were collected for 174 zircon grains from 2 rock

samples from ODP core 1270D, 5 rock samples from ODP core 1275D, and 11 rock

samples from the Atlantis Massif; 16 of the 18 rock samples have corresponding trace

element data (Table 1 and Table 2). Hafnium isotopes were first corrected for age so the

amount of radiogenic Hf ingrowth over time could be quantified. ɛHf at present time

[ɛ(0)] and ɛHf corrected for age [ɛ(t)] (age ranges from 0.48-2.45 Ma; Grimes et al.,

2011) differ by less than 0.1 epsilon value. The ingrowth of radiogenic Hf over the time since crystallization of the young (< 2.5 Ma) rocks is therefore negligible and does not modify ɛ(0) (Figure 20).

After data quality screening described in section 2.3, forty analyses (out of 174)

are omitted due to high isobaric interference correction, high internal error, and analytical

inconsistencies (Table 2). Remaining hafnium isotope ratios collected from the three

sample locations along the Mid-Atlantis Ridge span a broad range of ɛHf values from 56

13.3 to 22.1. The range of values for each sample, along with the arithmetic mean, is

summarized in Table 3. Most individual rocks exhibit within-sample variation from 1 to 4

epsilon units, high MSWD values, and span a wider range of values than the analytical

standard FC-1 (variation of 2.7 epsilon units). These observations are interpreted to

indicate derivation from a magma with heterogeneous hafnium rather than analytical

scatter about a single value.

Data are further analyzed using histograms, probability density functions, error-

weighted mean plots, and mean squared weighted deviates, which are compiled in

Appendix B. The weighted probability density function bases the probability that a data

point will fall within a given interval on the standard deviation of each data point (for this

case, the 2σ internal ɛHf errors for each analysis). The error-weighted mean is different

from an arithmetic mean in that all data points do not contribute equally to the average; it

( ) = σ 1/ is defined as: / ; where is the standard deviation, is the mean, is 1 2 2 2 휎 휇 ∑ 1 휎푖 휇 휎푖 the weighing factor (the inverse square of the error), and is the assigned error on each

푖 data point. The mean squared weighted deviate (MSWD)휎 is used to express scatter within

a dataset. If the scatter of data points is equivalent to that predicted from the analytical errors alone, then the MSWD will be near 1; excess scatter than predicted yield MSWD >

1 and less scatter than predicted yields MSWD <1. MSWD is defined as: 1

( ) 푁−1 ∗

2, where is the average values of , is the measured values, and is its 푁 푥푖−푥̅ 2 ∑푖=1 휎푥푖 푥̅ 푥푖 푥푖 휎푥푖 associated error. 57

Single grain ɛHf analyses from core 1270D range from 13.3 to 16.8 (Figure B1a).

Rock sample 1270D-19 has ɛHf values ranging from 13.3 to 16.8 (Figure B1b, Table 3,

Figure 21). Rock sample 1270D-25 has ɛHf values ranging from 13.3 to 16.4 ɛHf (Figure

B1c). A student’s t-test comparing the two samples from 1270D show that the ɛHf values in zircons from 19 and 25 mbsf are not statistically significant at 95% confidence. Zircon

ɛHf values measured at 1270D represent the lowest values in this study (Table 3; Figure

26).

Single zircon analyses from 1275D range from 15.4 to 22.1 ɛHf (Figure B2a). The shallowest sample, 1275D-10, contains the 12 highest ɛHf values in the core and in this study (Figure B2b, Table 3, Figure 21) ranging from 18.1 to 22.1; which a student’s t-test shows are statistically different from the other rock samples (134 – 200 mbsf) at p > 99%.

Samples 75D-134 and 75D-200 (Figures B2c, B2f) have MSWD values less than 1 (at

0.43 and 0.31, respectively); these are the only two samples from the 15˚20’ FZ that contain analyses that completely overlap at the 2 sigma level. A student’s t-tests show that ɛHf in zircons from 200 mbsf are statistically distinct from ɛHf in zircons from 134 and 180 mbsf at p < 0.05. Overall, these observations indicate that 1275D-10 is distinct from the other rock samples down-core, and small geochemical variations exist between the rock samples from 130 – 200 mbsf.

There are a total of ten rock samples from the Atlantis Massif: six from core

1309D and four samples collected along the surface of the southern wall and extend from

7.2 to 13 km away from the MAR axis (Table 3). All analyses from the Atlantis Massif display a range from 13.3 to 20.7 ɛHf (Figures B4a, 26); the 7.4 epsilon unit difference is 58

the largest range of values observed at any location in this study (Table 3). Analyses from

IODP core 1309D define the lowest and highest ɛHf values (13.3 – 20.7 ɛHf) for the

Atlantis Massif (Figure B3a). The range of values within individual rocks is small (2.3 –

4.4 epsilon units; Table 3), but in detail, small differences between rock samples are

indicated by visual inspection and supported by student’s t-tests. The ɛHf values from

sample 1309D-58 are statistically significant, at p < 0.05, from the rocks at 250, 564, 820,

1327, and 1415 mbsf. Further a t-test indicated ɛHf values from 250 mbsf are statistically

different from zircons at 820 mbsf for p < 0.05.

The four samples recovered on the surficial southern wall have values that range

from 13.7 to 19.1 ɛHf (Figure B4b). Surficial samples overlap the range defined by

samples from core 1309D (Figure B4c-f), and a student’s t-test state that ɛHf values from

1309D and the southern wall are not statistically different for p < 0.05. Similar statistical comparisons of all four rock samples from the southern wall of the Atlantis Massif indicate no statistical difference for p < 0.05.

Table 1. SHRIMP-RG Trace Element Results for Zircons from 15˚20' FZ and 30˚N, MAR Hf Sample ID TE Sample ID P Y Nb La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Th U ODP Hole 1270D: 70D-19, Highly Altered Dike Intruding Harzburgite a 70D_19_09, 09a 70D-19-2 189 682 6 0.0 12 0.0 0.6 1 0.6 13 5 59 26 125 30 255 52 10295 43 70 70D_19_12 70D-19-3 232 1633 3 0.0 10 0.1 1.7 4 2.0 42 15 177 69 321 68 523 99 8728 25 33 70D_19_13 70D-19-4 233 1714 4 0.0 9 0.0 2.1 4 2.2 42 15 176 72 322 70 530 102 7901 26 36 70D_19_13 70D-19-5 293 1370 9 0.0 9 0.0 0.8 2 1.1 23 9 123 56 270 61 520 102 7985 17 37 70D-19-6.1 424 1496 17 0.0 16 0.0 0.4 1 0.3 20 9 128 54 267 59 487 97 12925 55 147 ODP Hole 1270D: 70D-25, Highly Altered Dike Intruding Harzburgite a 70D-25-1.1 224 859 4 0.0 7 0.0 0.5 1 0.8 17 6 86 36 166 37 298 58 8543 12 23 70D-25-1.2 203 1428 3 0.0 9 0.0 0.9 3 1.5 31 11 137 54 237 50 372 69 7025 14 18 70D_25_05 70D-25-3 252 935 4 0.0 8 0.1 0.6 2 0.7 20 7 96 38 187 41 341 65 8871 14 25 70D_25_07 70D-25-4 190 1548 3 0.0 7 0.1 1.6 3 1.9 36 13 161 64 289 60 464 87 7853 19 26 70D_25_10 70D-25-5.1 202 755 3 0.0 6 0.0 0.5 1 0.7 16 6 79 31 146 33 264 50 8569 10 18 70D_25_10 70D-25-5.2 187 1552 3 0.0 7 0.1 1.5 4 1.8 34 13 149 61 267 54 420 78 6885 15 20 70D-25-6 214 1388 2 0.0 8 0.1 1.3 3 1.7 32 11 134 53 228 47 366 69 6999 15 19 ODP Hole 1275D: 75D-10, Brecciated Leucocratic Dikes 75D_10_3a, b 75D-10-3 379 1602 2 0.0 4 1.2 4 1.0 39 13 145 55 256 51 364 66 6559 7 12 75D_10_4 75D-10-4 294 816 1 0.0 3 0.4 1 0.4 16 6 81 32 157 34 273 49 9444 5 14 75D_10_6 75D-10-6 363 957 2 0.0 3 0.5 2 0.5 17 8 94 39 187 40 342 67 8706 6 17 75D_10_9 75D-10-9 554 1444 2 0.0 5 1.1 3 1.0 30 12 137 57 268 59 490 89 8351 24 43 75D_10_12 75D-10-12 284 762 2 0.0 3 0.4 1 0.5 16 6 73 31 147 32 259 46 9069 6 14 75D_10_13 75D-10-13 514 1604 3 0.0 6 1.1 3 1.0 31 12 143 61 290 61 498 89 8088 20 35 75D_10_14 75D-10-14 211 586 1 0.0 3 0.2 1 0.3 10 4 51 23 107 22 193 36 9357 4 11 75D_10_15 75D-10-15 370 2360 1 0.1 4 2.4 5 1.7 61 22 244 96 427 86 668 113 8380 19 30 75D_10_16 75D-10-16 297 788 1 0.0 4 0.4 2 0.5 16 6 73 31 142 32 262 47 9019 6 16 75D-10-18 523 3443 2 0.1 5 2.9 7 2.5 90 33 366 142 629 125 965 165 8433 33 47 75D_10_19 75D-10-19 361 979 1 0.0 4 0.8 2 0.6 18 8 90 37 180 38 324 59 8680 9 21 ODP Hole 1275D: 75D-134, Leucocratic Dike Intruding Gabbro 75D_134_01 75D-134-1 238 1486 2 0.0 4 0.7 3 0.9 33 13 155 61 279 56 464 82 9627 16 31 75D_134_01 75D-134-1.1 251 767 2 0.0 4 0.4 1 0.3 13 6 72 31 152 32 272 51 10016 13 37 75D_134_02 75D-134-2 289 796 1 0.0 2 0.4 1 0.4 14 6 77 32 151 33 279 52 8522 7 18 75D_134_03, a 75D-134-3 461 2724 4 0.0 5 2.0 6 2.1 66 25 286 114 524 105 837 145 8223 40 61 75D_134_04 75D-134-4 265 766 2 0.0 2 0.4 1 0.5 16 6 75 32 143 30 256 45 8462 7 17 75D_134_05 75D-134-5 95 714 3 0.0 2 0.4 1 0.5 16 6 69 35 149 34 294 54 11720 11 26 75D_134_06 75D-134-6 366 2322 2 0.0 4 2.3 5 1.6 54 20 237 98 427 85 687 119 8290 28 46 75D_134_07 75D-134-7 238 654 2 0.0 3 0.3 1 0.4 12 5 62 26 126 26 225 41 9041 7 20 75D_134_08 75D-134-8 241 708 2 0.0 3 0.4 1 0.4 13 6 70 29 136 30 254 46 9201 8 22 75D-134-9 294 877 2 0.0 3 0.4 1 0.5 17 7 81 36 167 37 311 58 8532 8 21 ODP Hole 1275D: 75D-144, Leucocratic Dike Intruding Gabbro 75D_144_06 75D-144-6 417 2729 3 0.0 9 2.7 6 1.3 57 22 268 113 508 107 839 145 9903 57 95 75D_144_08 75D-144-8 286 1750 2 0.0 5 2.2 5 1.5 46 16 188 77 344 69 562 98 8995 21 35 75D_144_09 75D-144-9 293 1690 2 0.0 5 2.2 4 1.4 44 16 178 72 323 66 525 95 8992 20 34 75D_144_15 75D-144-15 294 1679 2 0.0 5 1.5 5 1.3 45 16 177 70 308 66 516 90 9593 20 35 75D_144_18 75D-144-18 304 1523 1 0.0 4 2.3 5 1.3 42 14 159 64 286 57 456 80 8900 12 23 75D-144-21 313 2056 2 0.0 5 2.5 5 1.5 44 17 200 84 378 77 624 107 8764 24 41 75D-144-23 329 1949 2 0.0 5 2.6 5 1.6 45 17 201 81 365 75 593 104 8917 23 39 75D-144-26 315 2007 2 0.0 6 2.3 5 1.4 47 17 198 84 391 79 635 111 9443 32 54 75D-144-27 300 1675 1 0.0 5 1.7 4 1.3 41 15 173 68 310 64 496 88 9410 19 34 75D-144-29 315 1713 1 0.0 6 1.5 4 1.3 40 15 176 70 320 64 519 93 9535 20 37 ODP Hole 1275D: 75D-180, Leucocratic Dike Intruding Gabbro 75D_180_1 75D-180-1 432 2708 3 0.0 8 2.4 6 1.6 59 23 269 111 512 103 811 143 9530 48 79 75D_180_2 75D-180-2 414 2333 3 0.0 7 2.4 6 1.4 53 20 235 98 440 91 707 125 9340 38 62 75D_180_3 75D-180-3 337 1068 4 0.0 5 0.6 2 0.5 21 8 99 43 203 45 380 72 9823 18 44 60

Table 1. (continued) Hf Sample ID TE Sample ID P Y Nb La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Th U 75D_180_4 75D-180-4 408 2518 3 0.0 7 2.5 6 1.5 62 23 260 107 466 95 763 129 9405 42 66 75D_180_6 75D-180-6 375 901 2 0.0 4 0.6 2 0.5 19 7 89 36 179 37 317 60 8435 9 21 75D_180_7 75D-180-7 416 2673 4 0.0 8 2.7 6 1.4 56 22 264 111 507 104 823 147 9449 52 84 75D_180_8 75D-180-8 383 2186 2 0.0 7 2.4 5 1.3 48 18 212 88 408 81 644 115 9742 38 66 75D_180_12 75D-180-12 435 903 2 0.0 4 0.6 1 0.6 17 7 85 37 179 38 317 59 8734 12 29 75D_180_14 75D-180-14 335 1977 2 0.0 7 1.7 5 1.3 49 17 204 81 365 75 604 104 9627 31 51 75D_180_16 75D-180-16 464 2954 4 0.0 8 3.1 6 1.6 67 24 287 122 531 109 860 150 9465 63 96 75D-180-18 444 2559 2 0.0 7 2.9 6 1.6 61 22 259 107 466 96 760 129 8939 39 63 75D-180-21 547 3176 4 0.0 8 3.2 7 1.8 73 29 332 134 592 121 948 163 9226 57 85 ODP Hole 1275D: 75D-200, Leucocratic Dike Intruding Gabbro 75D_200_1 75D-200-1 562 3641 4 0.0 6 3.2 9 3.2 99 34 389 152 647 133 1053 179 7871 55 69 75D_200_2 75D-200-2 455 1491 4 0.0 5 0.9 2 0.8 25 10 130 58 300 66 574 107 8221 22 57 75D_200_5 75D-200-5 376 1084 2 0.0 5 0.9 2 0.7 22 9 101 45 203 42 358 66 8339 20 38 75D_200_7 75D-200-7 304 895 2 0.0 5 0.5 1 0.4 16 7 86 35 168 38 313 58 9401 15 37 75D_200_10 75D-200-10 264 1557 1 0.0 3 1.3 4 1.3 37 13 147 60 257 53 423 74 8816 14 25 75D_200_11 75D-200-11 236 1262 1 0.0 3 0.6 2 0.8 28 10 132 53 232 49 387 67 9265 15 27 75D_200_12 75D-200-12 274 793 2 0.0 3 0.3 1 0.4 14 6 76 31 153 32 276 48 9810 8 20 IODP Hole U1309D: 1309D-564, Fe-Ti Oxide Gabbro a D564-7 0.0 10 0.7 3 0.9 37 222 517 1009 11596 125 206 1309D_564_16 D564-5 0.0 3 1.2 3 0.9 34 168 346 595 15004 181 172 1309D_564_09 D564-4 0.0 7 3.2 9 2.2 91 376 632 948 9486 47 76 1309D_564_04 D564-2 0.0 20 2.7 8 0.7 86 380 683 996 13244 336 440 1309D_564_01 D564-1 0.1 6 3.7 7 2.4 69 289 495 759 9309 36 63 IODP Hole U1309D: 1309D-820, Fe-Ti Oxide Gabbro a 1309D_820_01 D820-1 0.0 44 2.5 11 3.7 149 660 1045 1546 9318 77 134 1309D_820_02 D820-2 0.1 64 11.5 31 5.5 312 1235 1774 2280 12672 943 636 1309D_820_03 D820-3 0.0 71 4.1 16 3.1 207 997 1665 2409 12399 1102 896 1309D_820_04 D820-4 0.0 11 1.4 4 0.8 49 218 375 577 8135 38 52 1309D_820_5, a D820-5 0.0 34 1.9 9 2.8 120 576 937 1414 15552 1792 821 D820-6 0.0 108 5.0 21 5.4 279 1481 2726 4158 14613 1272 600 1309D_820_07 D820-7 0.2 20 3.1 9 1.4 104 471 816 1195 12214 595 315 1309D_820_08 D820-8 0.1 4 1.6 5 1.8 54 233 421 695 11831 10 21 D820-9 0.1 96 16.2 50 15.0 514 1933 2595 3332 9435 117 174 IODP Hole U1309D: 1309D-1175, Fe-Ti Oxide Gabbro a 1309D_1175_01 D1175-15 0.1 162 8.0 20 3.5 218 992 1883 3472 19368 174 107 1309D_1175_02y D1175-14 0.0 50 2.6 9 1.4 113 584 1163 2263 16871 230 153 1309D_1175_06 D1175-12 0.1 163 8.6 38 6.5 477 2335 3488 5558 14086 269 321 1309D_1175_07y D1175-11 0.0 32 3.6 10 1.0 114 424 630 831 5760 319 53 D1175-10 1.0 393 8.1 28 5.1 383 1904 3900 7004 9892 2884 1102 1309D_1175_10y D1175-9 0.1 597 18.8 56 8.3 568 2292 3253 4604 12506 263 73 1309D_1175_13y D1175-8 0.5 787 21.6 51 7.4 495 1730 2785 4722 12638 534 225 1309D_1175_15 D1175-7 0.1 519 15.7 38 6.0 364 1270 2074 3856 16742 241 89 1309D_1175_17 D1175-6 0.0 265 4.7 20 5.0 279 1676 3926 7065 12262 3908 1616 1309D_1175_20 D1175-5 0.1 68 16.7 45 10.1 446 1721 2377 3007 9198 70 96 D1175-4 0.3 159 30.3 74 11.3 689 2980 3475 8777 223 191 D1175-3 0.2 211 16.8 50 5.3 511 1912 2721 3527 13075 475 247 D1175-2 0.1 246 4.4 20 4.1 271 1448 2640 3895 10651 198 262 1309D_1175_25y D1175-1 0.1 35 6.1 17 2.8 168 794 1456 2643 11557 99 68 IODP Hole U1309D: 1309D-1327, Fe-Ti Oxide Gabbro a 1309D_1327_01 D1327-9 0.0 23 3.1 13 4.0 130 704 1089 1594 6269 9 133 1309D_1327_02 D1327-8 0.0 14 2.6 7 2.1 68 314 593 990 9093 31 45 1309D_1327_03 D1327-7 0.1 43 6.8 16 4.5 165 806 1542 2519 9989 168 174 1309D_1327_04 D1327-6 0.0 13 0.8 3 1.1 31 157 298 513 14255 64 99 61

Table 1. (continued) Hf Sample ID TE Sample ID P Y Nb La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Th U D1327-5 1.8 119 42.2 68 2.9 514 1966 2929 3891 15518 450 1532 1309D_1327_06 D1327-4 0.1 39 6.5 15 2.9 139 546 801 1058 10861 275 147 1309D_1327_07 D1327-3 0.0 24 2.4 7 1.6 75 337 600 952 10351 29 44 1309D_1327_09 D1327-2 0.1 16 3.5 9 1.8 86 360 630 982 9881 25 31 1309D_1327_10 D1327-1 0.0 10 1.3 4 1.4 43 202 377 645 9700 12 21 IODP Hole U1309D: 1309D-1415, Leucocratic Melt Intrusions a 1309D_1415_25 D1415-12.2 1038 5043 5 0.0 47 0.1 1.9 7 2.3 85 34 438 192 922 203 1686 295 11165 35 71 1309D_1415_24 D1415-12.1 5263 8794 49 1.4 190 0.4 6.3 15 5.0 147 54 670 315 1455 341 3126 527 13215 7556 3106 1309D_1415_22 D1415-10.1 1455 19897 76 0.2 1271 1.0 22.3 67 19.5 631 223 2450 1012 4028 742 5359 811 8494 1945 1883 D1415-8r 2384 350 1 0.0 3 0.0 0.1 0 0.1 3 1 21 11 67 19 210 49 14690 44 106 1309D_1415_15 D1415-7.1 1039 6054 5 0.0 58 0.1 2.4 9 3.1 109 43 540 237 1078 233 1898 330 10012 58 96 1309D_1415_15 D1415-7.2 3082 3458 1 0.0 23 0.1 1.6 6 2.2 68 26 316 134 608 129 1054 182 9859 18 31 D1415-6.1 1726 4038 3 0.0 115 0.1 3.7 10 2.1 92 32 387 167 803 183 1571 262 6241 14 40 D1415-5r 1790 703 1 0.0 4 0.0 0.0 0 0.2 5 3 43 24 139 39 415 99 11694 93 184 D1415-2.1 353 20497 10 0.5 368 1.1 23.3 47 12.1 342 127 1499 783 3014 694 6299 1046 10006 411 377 1309D_1415_01 D1415-1 169 3949 1 0.0 16 0.1 2.9 8 3.0 83 31 375 159 700 147 1186 207 10431 17 29 Alvin Dive 3646-1205: Talc-Amphibolite Schist a 3646_1205_02 3646-1205-2 601 2171 3 0.0 9 0.1 0.7 2 1.0 33 14 187 80 386 84 677 128 9783 5 13 3646_1205_3a, b 3646-1205-3 726 3265 3 0.0 15 0.1 1.6 5 2.1 67 26 343 132 634 133 1054 191 11200 15 27 3646_1205_6, a 3646-1205-6 369 1375 2 0.0 10 0.1 0.4 2 0.5 20 9 123 55 268 60 497 94 13542 8 17 3646_1205_07 3646-1205-7 514 2068 3 0.0 15 0.1 0.9 3 1.2 36 15 200 82 394 86 683 126 11827 60 126 3646_1205_08 3646-1205-8 608 2257 5 0.6 19 0.2 2.3 4 1.4 43 18 232 96 440 92 726 134 11549 10 17 Alvin Dive 3652-1333: Foliated Tremolite+Chlorite-Rich Rock a 3652-1333-1.1 3195 23503 175 0.1 201 0.1 5.6 19 1.9 306 139 1940 818 3860 818 5978 869 21487 2931 1765 3652-1333-3.1 2403 17137 45 0.1 189 0.2 9.0 30 5.3 401 156 1884 717 3028 581 4025 617 9938 388 414 3652_1333_06 3652-1333-4.1 1044 6321 11 0.0 42 0.1 2.4 8 1.9 122 50 646 261 1170 235 1740 292 11109 78 152 3652_1333_07 3652-1333-5.1 2356 17362 47 0.1 212 0.2 10.5 34 5.6 429 160 1886 724 2940 548 3726 570 9464 584 525 3652_1333_07 3652-1333-5.2 2639 19027 48 0.1 246 0.2 10.5 35 4.9 441 173 2037 779 3162 602 4122 618 9797 633 568 3652_1333_07 3652-1333-5.3 2720 19065 48 0.1 287 0.2 10.9 37 4.4 478 181 2095 789 3222 600 4083 615 10999 984 718 3652_1333_10 3652-1333-6.2 1187 8079 14 0.0 96 0.1 2.3 8 1.0 120 53 731 314 1423 301 2184 356 19777 298 442 Alvin Dive 3647-1359: Breciated Amphibolite-Chlorite Rock a 3647_1359_02 3647-1359-2 275 1028 1 0.0 3 0.2 0.4 2 0.7 21 8 110 42 193 42 327 61 9633 2 5 3647_1359_03 3647-1359-4 2902 19193 17 0.1 68 0.3 10.4 35 8.9 464 176 2179 826 3439 687 5132 880 14408 2163 1202 3647_1359_05 3647-1359-5 178 616 1 0.0 3 0.1 0.3 1 0.4 11 5 61 25 122 28 219 43 11274 3 6 3647_1359_06 3647-1359-6.1 805 6421 23 0.1 38 0.2 3.7 13 4.2 156 60 750 286 1202 231 1588 262 10327 98 97 3647_1359_06 3647-1359-6.2 920 6829 46 0.0 50 0.1 2.5 10 2.7 143 59 773 306 1336 270 1947 326 12205 83 140 3647-1359-7 403 2605 1 0.0 5 0.1 1.4 4 1.7 52 20 265 108 501 109 872 167 15161 49 77 Dredge D3-21: Metasomatized Amphibolite+Chlorite+Talc-Rich Fault Rock a D3_21_04 D3-21-1 360 1409 1 0.0 7 0.1 0.7 2 0.8 27 11 147 58 283 62 487 92 10714 5 10 D3_21_05 D3-21-2 241 949 1 0.0 5 0.1 0.4 1 0.5 16 7 97 40 196 44 351 68 12200 4 10 D3_21_06 D3-21-3 316 1839 1 0.2 5 0.1 1.3 3 1.8 42 16 200 80 354 73 540 99 10585 5 10 D3_21_07 D3-21-4 1031 5078 8 0.1 23 0.1 3.8 10 2.2 128 48 558 207 856 165 1138 191 12684 49 69 D3_21_11 D3-21-5 969 5293 7 0.2 22 0.1 4.1 11 2.7 135 51 584 218 891 171 1165 195 12619 50 69 D3-21-7 492 2683 4 0.0 8 0.1 2.1 6 1.8 65 24 284 107 475 94 689 124 13613 24 44 SHRIMP-RG Trace Element analyses were conducted at Stanford University. All Trace Element values are expressed in parts per million (ppm).

a Data from Grimes et al., 2009.

REE Averages ODP 1270D 100000. 10000. e e t t i i r r 10000. 1000. d d n n 1000. o o 100. C h

C h 100. / / t

t 10. n n 10. ODP 1270D (avg) e e 1.0 1270D-19 (avg) ODP 1275D (avg) m

m 1.0 e e l l Atlantis Massif (avg) 1270D-25 (avg) E E 0.10 0.10 0.01 0.01 La Ce Pr NdSmEu Gd Tb DyHo Er TmYb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

ODP 1275D Atlantis Massif 10000. 100000. e e t t i i r

1000. r 10000. d d n

n 1000. o

100. o 100. C h 1275D-10 (avg) C h 1309D-564 (avg) / 10. / 3646-1205 (avg) t 1275D-134 (avg) t 10. 1309D-820 (avg) n n 3652-1333 (avg) e 1.0 1275D-144 (avg) e 1.0 1309D-1175 (avg) m

m 3647-1359 (avg)

e 1275D-180 (avg)

e 1309D-1327 (avg) l 0.10 l D3-21 (avg)

E 0.10 1275D-200 (avg) E 1309D-1415 (avg) 0.01 0.01 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd SmEu Gd Tb Dy Ho Er TmYb Lu

Figure 18: Rare earth element diagrams for zircons from cores at ODP 1270D, ODP 1275D, and Atlantis Massif along the Mid-Atlantic Ridge. Normalizing chondrite values are from Anders and Grevasse (1989) multiplied by 1.3596 Korotev (1996) Wed Site. Wash. U. (REE diagrams for individual samples are provided in Appendix A). 63

A. 100

Fractionation

b Source Enrichment Y

/ 1

U Enrichment

0.1

0.01 10 100 1000 10000 100000 Y (ppm)

B. 100

Source Enrichment b Y

/ 1 U

Fractionation Enrichment 0.1

0.01 5000 50000 Hf (ppm)

ODP 1270D ODP 1275D Atlantis Massif

Figure 19: Trace element bivariate plots for U/Yb plotted against A. Y and B. Hf. Black arrows indicate expected general trends for source enrichment (as for enriched mantle sources and continental crustal settings) and crystal fractionation of minerals with low REE concentrations. Data from this study and Grimes et al. (2009). 64

Table 2. LA-ICP-MS Lu-Hf Isotope Analysis of Zircons from 15˚20' FZ and 30˚N, MAR Sample Name 176Lu/177Hf 1 σ errora 176Hf/177Hf 1 σ errora ɛ(0)b 2 σ error % Corrc ODP Hole 1270D: 70D-19, Highly Altered Dike Intruding Harzburgite 70D_19_01 0.00204 0.00002 0.28317 0.00000 13.5 0.3 25% 70D_19_02 0.00138 0.00001 0.28318 0.00006 14.0 4.5 18% 70D_19_03 0.00066 0.00001 0.28326 0.00000 16.8 0.1 11% 70D_19_03a 0.00048 0.00001 0.28325 0.00001 16.5 0.7 9% 70D_19_04 0.00152 0.00002 0.28325 0.00001 16.4 0.8 19% 70D_19_05 0.00116 0.00001 0.28317 0.00001 13.6 0.4 15% 70D_19_06 0.00147 0.00001 0.28316 0.00001 13.3 1.0 18% 70D_19_09 0.00387 0.00002 0.28323 0.00024 15.6 17.0 41% 70D_19_09a 0.00182 0.00003 0.28309 0.00003 10.9 2.3 17% 70D_19_12 0.00185 0.00002 0.28320 0.00001 14.6 0.6 23% 70D_19_13 0.00170 0.00001 0.28319 0.00002 14.3 1.1 21% 70D_19_15 0.00497 0.00002 0.28309 0.00044 10.9 31.1 46% ODP Hole 1270D: 70D-25, Highly Altered Dike Intruding Harzburgite 70D_25_05 0.00102 0.00001 0.28320 0.00001 14.6 0.4 14% 70D_25_07 0.00073 0.00001 0.28319 0.00005 14.3 3.3 11% 70D_25_10 0.00118 0.00002 0.28318 0.00002 14.1 1.2 15% 70D_25_11 0.00167 0.00002 0.28316 0.00003 13.3 2.1 21% 70D_25_15 0.00200 0.00002 0.28319 0.00007 14.3 4.9 25% 70D_25_20 0.00121 0.00002 0.28325 0.00002 16.4 1.1 16% 70D_25_20a 0.00312 0.00002 0.28314 0.00014 12.4 9.9 36% 70D_25_21 0.00157 0.00001 0.28319 0.00001 14.1 0.5 20% ODP Hole 1275D: 75D-10, Brecciated Plagiogranite 75D_10_3a 0.00147 0.00001 0.28339 0.00004 21.4 2.9 19% 75D_10_3b 0.00071 0.00001 0.28334 0.00000 19.6 0.1 10% 75D_10_4 0.00081 0.00002 0.28334 0.00001 19.4 0.4 12% 75D_10_5 0.00070 0.00001 0.28337 0.00001 20.6 0.5 10% 75D_10_6 0.00090 0.00002 0.28334 0.00002 19.5 1.1 13% 75D_10_9 0.00120 0.00002 0.28341 0.00005 22.1 3.2 17% 75D_10_12 0.00088 0.00001 0.28337 0.00000 20.7 0.3 13% 75D_10_13 0.00129 0.00002 0.28332 0.00001 19.0 0.5 18% 75D_10_14 0.00068 0.00002 0.28339 0.00002 21.5 1.7 10% 75D_10_15 0.00249 0.00002 0.28330 0.00002 18.1 1.6 31% 75D_10_16 0.00080 0.00002 0.28330 0.00000 18.4 0.2 12% 75D_10_19 0.00114 0.00002 0.28338 0.00003 21.0 2.0 16% 75D_10_21 0.00088 0.00001 0.28335 0.00002 20.0 1.6 12% 75D_10_22 0.00228 0.00001 0.28336 0.00003 20.3 2.3 28% 75D_10_24 0.00235 0.00002 0.28336 0.00002 20.2 1.1 28% ODP Hole 1275D: 75D-134, Plagiogranite Dike Intruding Gabbro 75D_134_01 0.00143 0.00002 0.28329 0.00007 17.7 4.8 20% 75D_134_02 0.00089 0.00001 0.28328 0.00001 17.5 0.6 13% 75D_134_03 0.00235 0.00002 0.28328 0.00006 17.5 4.5 29% 75D_134_03a 0.00079 0.00002 0.28328 0.00000 17.6 0.1 12% 75D_134_04 0.00115 0.00002 0.28326 0.00002 16.7 1.6 16% 75D_134_05 0.00133 0.00001 0.28328 0.00003 17.6 2.1 18% 75D_134_06 0.00117 0.00001 0.28328 0.00001 17.5 0.5 16% 75D_134_07 0.00129 0.00001 0.28328 0.00009 17.6 6.6 17% 75D_134_08 0.00071 0.00002 0.28327 0.00001 17.3 0.4 10% ODP Hole 1275D: 75D-144, Plagiogranite Dike Intruding Oxide-Bearing Gabbro 75D_144_06 0.00149 0.00002 0.28333 0.00005 19.2 3.5 20% 75D_144_08 0.00166 0.00001 0.28329 0.00003 18.0 1.8 21% 65

Table 2. (continued) Sample Name 176Lu/177Hf 1 σ errora 176Hf/177Hf 1 σ errora ɛ(0)b 2 σ error % Corrc 75D_144_09 0.00169 0.00001 0.28330 0.00004 18.0 2.5 22% 75D_144_15 0.00158 0.00002 0.28324 0.00000 16.1 0.2 21% 75D_144_18 0.00148 0.00002 0.28325 0.00010 16.6 6.7 19% ODP Hole 1275D: 75D-180, Plagiogranite Dike Intruding Gabbro 75D_180_1 0.00188 0.00001 0.28327 0.00003 17.0 1.8 25% 75D_180_2 0.00241 0.00002 0.28327 0.00000 17.1 0.3 30% 75D_180_3 0.00125 0.00001 0.28329 0.00003 18.0 1.8 17% 75D_180_4 0.00265 0.00002 0.28329 0.00004 17.8 2.8 32% 75D_180_6 0.00116 0.00001 0.28332 0.00000 18.8 0.0 16% 75D_180_7 0.00265 0.00002 0.28329 0.00002 17.9 1.1 32% 75D_180_8 0.00244 0.00002 0.28328 0.00003 17.4 2.0 30% 75D_180_12 0.00118 0.00002 0.28330 0.00001 18.3 0.7 17% 75D_180_14 0.00217 0.00002 0.28323 0.00005 15.7 3.5 27% 75D_180_16 0.00278 0.00002 0.28322 0.00000 15.4 0.3 34% ODP Hole 1275D: 75D-200, Plagiogranite Dike Intruding Gabbro 75D_200_1 0.00182 0.00001 0.28332 0.00010 19.1 6.8 22% 75D_200_2 0.00238 0.00001 0.28331 0.00005 18.6 3.6 27% 75D_200_3 0.00110 0.00001 0.28328 0.00003 17.6 2.4 15% 75D_200_4 0.00170 0.00001 0.28329 0.00007 17.8 5.2 21% 75D_200_5 0.00104 0.00002 0.28330 0.00001 18.0 0.8 14% 75D_200_7 0.00093 0.00001 0.28332 0.00002 19.0 1.6 13% 75D_200_10 0.00188 0.00002 0.28330 0.00015 18.3 10.6 23% 75D_200_11 0.00257 0.00001 0.28331 0.00022 18.7 15.6 29% 75D_200_12 0.00085 0.00001 0.28331 0.00002 18.4 1.1 12% IODP Hole U1309D: 1309D-58, Leucocratic Dike Intruding Gabbro 1309D_58_02 0.00411 0.00002 0.28313 0.00003 12.3 2.0 42% 1309D_58_12 0.00345 0.00001 0.28319 0.00009 14.2 6.1 38% 1309D_58_16 0.00299 0.00001 0.28319 0.00002 14.2 1.6 33% 1309D_58_19 0.00415 0.00002 0.28315 0.00001 12.8 0.9 42% 1309D_58_20 0.00378 0.00002 0.28319 0.00008 14.2 5.9 39% 1309D_58_21 0.00379 0.00002 0.28323 0.00002 15.6 1.1 39% 1309D_58_24 0.00343 0.00002 0.28318 0.00001 14.0 0.4 37% 1309D_58_32 0.00391 0.00002 0.28320 0.00003 14.6 1.8 40% 1309D_58_34 0.00325 0.00002 0.28316 0.00002 13.3 1.1 35% IODP Hole U1309D: 1309D-250, Fe-Ti Oxide Gabbro 1309D_250_03 0.00136 0.00002 0.28328 0.00000 17.4 0.1 19% 1309D_250_07 0.00226 0.00002 0.28320 0.00000 14.6 0.3 28% 1309D_250_08 0.00428 0.00002 0.28317 0.00002 13.7 1.3 43% 1309D_250_14 0.00278 0.00002 0.28322 0.00000 15.5 0.2 33% 1309D_250_15 0.00169 0.00001 0.28321 0.00000 15.1 0.3 23% 1309D_250_16 0.00169 0.00002 0.28324 0.00001 16.2 0.8 23% 1309D_250_17 0.00164 0.00002 0.28322 0.00000 15.5 0.1 23% 1309D_250_19 0.00120 0.00002 0.28325 0.00000 16.5 0.1 17% 1309D_250_21 0.00169 0.00002 0.28326 0.00000 16.7 0.1 22% 1309D_250_24 0.00172 0.00002 0.28325 0.00006 16.4 4.1 22% IODP Hole U1309D: 1309D-564, Fe-Ti Oxide Gabbro 1309D_564_01 0.00151 0.00002 0.28333 0.00002 19.3 1.7 22% 1309D_564_03 0.00096 0.00002 0.28327 0.00003 17.2 1.9 14% 1309D_564_04 0.00150 0.00001 0.28330 0.00001 18.2 0.4 21% 1309D_564_07 0.00579 0.00002 0.28341 0.00012 22.1 8.5 50% 1309D_564_09 0.00158 0.00001 0.28326 0.00003 16.6 2.1 22% 66

Table 2. (continued) Sample Name 176Lu/177Hf 1 σ errora 176Hf/177Hf 1 σ errora ɛ(0)b 2 σ error % Corrc 1309D_564_12 0.00265 0.00002 0.28331 0.00004 18.7 3.1 32% 1309D_564_12a 0.00154 0.00002 0.28322 0.00000 15.2 0.3 22% 1309D_564_16 0.00082 0.00001 0.28325 0.00000 16.4 0.2 12% 1309D_564_18 0.00164 0.00001 0.28325 0.00005 16.5 3.5 23% 1309D_564_19 0.00074 0.00001 0.28322 0.00004 15.3 2.5 12% 1309D_564_25 0.00046 0.00001 0.28325 0.00002 16.3 1.2 7% IODP Hole U1309D: 1309D-820, Fe-Ti Oxide Gabbro 1309D_820_01 0.00478 0.00002 0.28343 0.00005 22.8 3.4 50% 1309D_820_02 0.00324 0.00002 0.28332 0.00001 19.0 1.0 39% 1309D_820_03 0.00438 0.00002 0.28318 0.00004 13.9 2.6 46% 1309D_820_04 0.00236 0.00002 0.28337 0.00001 20.7 0.8 30% 1309D_820_05 0.00414 0.00002 0.28328 0.00002 17.4 1.2 44% 1309D_820_05a 0.00453 0.00004 0.28319 0.00024 14.4 17.0 38% 1309D_820_07 0.00333 0.00002 0.28332 0.00001 18.8 0.7 40% 1309D_820_08 0.00173 0.00001 0.28330 0.00001 18.1 0.8 24% IODP Hole U1309D: 1309D-1175, Fe-Ti Oxide Gabbro 1309D_1175_01 0.00384 0.00002 0.28319 0.00024 14.2 17.0 37% 1309D_1175_02y 0.00508 0.00002 0.28311 0.00018 11.6 12.7 48% 1309D_1175_03 0.00409 0.00001 0.28312 0.00015 11.9 10.6 40% 1309D_1175_06 0.00988 0.00003 0.28338 0.00008 20.9 5.3 66% 1309D_1175_07y 0.00374 0.00002 0.28348 0.00017 24.6 12.0 43% 1309D_1175_10y 0.01008 0.00003 0.28250 0.00042 -10.1 29.7 68% 1309D_1175_13y 0.01142 0.00004 0.28287 0.00016 3.0 11.3 67% 1309D_1175_15 0.00847 0.00003 0.28315 0.00013 12.9 9.2 59% 1309D_1175_17 0.01127 0.00003 0.28274 0.00020 -1.6 14.1 69% 1309D_1175_20 0.00806 0.00003 0.28314 0.00002 12.6 1.1 63% 1309D_1175_25y 0.00715 0.00002 0.28308 0.00009 10.4 6.2 55% IODP Hole U1309D: 1309D-1327, Fe-Ti Oxide Gabbro 1309D_1327_01 0.00431 0.00003 0.28356 0.00004 27.5 2.8 47% 1309D_1327_02 0.00235 0.00001 0.28326 0.00005 16.8 3.3 31% 1309D_1327_03 0.00707 0.00003 0.28353 0.00007 26.3 4.7 57% 1309D_1327_04 0.00233 0.00002 0.28322 0.00004 15.3 3.0 31% 1309D_1327_06 0.00144 0.00001 0.28320 0.00001 14.8 0.9 27% 1309D_1327_07 0.00305 0.00002 0.28333 0.00005 19.2 3.2 37% 1309D_1327_09 0.00247 0.00001 0.28329 0.00003 17.9 2.0 31% 1309D_1327_10 0.00122 0.00002 0.28329 0.00001 17.9 0.4 18% IODP Hole U1309D: 1309D-1415, Leucocratic Melt Intrusions 1309D_1415_01 0.00231 0.00001 0.28320 0.00004 14.8 2.7 29% 1309D_1415_10 0.00336 0.00001 0.28330 0.00008 18.2 5.9 35% 1309D_1415_15 0.00355 0.00002 0.28332 0.00004 19.0 3.1 38% 1309D_1415_18 0.00756 0.00003 0.28340 0.00006 21.9 4.2 57% 1309D_1415_20 0.00260 0.00002 0.28329 0.00013 17.9 9.2 29% 1309D_1415_21 0.00864 0.00003 0.28309 0.00021 10.9 14.9 61% 1309D_1415_22 0.01662 0.00021 0.28384 0.00068 37.3 48.1 76% 1309D_1415_23 0.01181 0.00005 0.28417 0.00022 49.0 15.6 70% 1309D_1415_24 0.03648 0.00012 0.28314 0.00030 12.4 21.2 86% 1309D_1415_25 0.00367 0.00002 0.28326 0.00002 16.9 1.6 39% 1309D_1415_27 0.00713 0.00002 0.28326 0.00003 16.6 2.1 55% Alvin Dive 3646-1205: Talc-Amphibolite Schist 3646_1205_01 0.00199 0.00002 0.28322 0.00000 15.3 0.3 26% 3646_1205_02 0.00210 0.00001 0.28322 0.00007 15.5 5.2 27% 67

Table 2. (continued) Sample Name 176Lu/177Hf 1 σ errora 176Hf/177Hf 1 σ errora ɛ(0)b 2 σ error % Corrc 3646_1205_03a 0.00176 0.00001 0.28321 0.00002 15.0 1.6 23% 3646_1205_03b 0.00123 0.00002 0.28322 0.00003 15.3 2.2 18% 3646_1205_04 0.00152 0.00002 0.28326 0.00003 16.7 1.8 21% 3646_1205_05 0.00206 0.00002 0.28321 0.00001 15.1 0.8 26% 3646_1205_06 0.00162 0.00001 0.28322 0.00000 15.3 0.2 22% 3646_1205_06a 0.00258 0.00001 0.28318 0.00001 14.0 0.5 32% 3646_1205_07 0.00224 0.00002 0.28317 0.00004 13.7 2.8 28% 3646_1205_08 0.00275 0.00002 0.28321 0.00005 14.9 3.2 32% Alvin Dive 3652-1333: Foliated Tremolite+Chlorite-Rich Rock 3652_1333_03 0.00621 0.00002 0.28316 0.00012 13.4 8.5 53% 3652_1333_04 0.00274 0.00002 0.28322 0.00003 15.5 1.8 33% 3652_1333_06 0.00383 0.00002 0.28326 0.00001 16.7 0.8 41% 3652_1333_07 0.00836 0.00006 0.28311 0.00009 11.6 6.2 63% 3652_1333_09 0.00218 0.00002 0.28330 0.00003 18.2 1.9 27% 3652_1333_10 0.00400 0.00002 0.28313 0.00014 12.3 9.9 41% 3652_1333_12 0.00142 0.00002 0.28327 0.00004 17.3 2.6 19% 3652_1333_13 0.00229 0.00002 0.28318 0.00001 14.0 0.7 29% 3652_1333_14 0.00349 0.00002 0.28318 0.00007 13.8 5.2 40% 3652_1333_20 0.00266 0.00002 0.28320 0.00000 14.6 0.3 32% Alvin Dive 3647-1359: Brecciated Amphibole-Chlorite Rock 3647_1359_02 0.00078 0.00001 0.28326 0.00001 16.6 0.4 11% 3647_1359_03 0.00080 0.00001 0.28326 0.00003 16.9 2.4 12% 3647_1359_05 0.00104 0.00002 0.28332 0.00003 19.1 2.1 14% 3647_1359_06 0.00102 0.00001 0.28323 0.00001 15.7 0.7 14% 3647_1359_08 0.00062 0.00002 0.28328 0.00001 17.4 0.6 10% 3647_1359_09 0.00107 0.00001 0.28325 0.00001 16.4 0.5 15% 3647_1359_11 0.00106 0.00002 0.28325 0.00003 16.5 2.1 15% Dredge D3-21: Metasomatized Amphibolite+Chlorite+Talc-Rich Fault Rock D3_21_01 0.00100 0.00001 0.28322 0.00001 15.5 0.5 14% D3_21_03 0.00116 0.00002 0.28325 0.00001 16.6 0.4 16% D3_21_04 0.00131 0.00001 0.28324 0.00001 16.1 0.4 18% D3_21_05 0.00115 0.00001 0.28321 0.00000 15.0 0.0 16% D3_21_06 0.00122 0.00002 0.28324 0.00000 16.1 0.2 17% D3_21_07 0.00119 0.00001 0.28332 0.00004 18.9 3.1 16% D3_21_11 0.00111 0.00001 0.28327 0.00000 17.0 0.2 15% D3_21_13 0.00119 0.00001 0.28326 0.00001 16.8 1.0 16% D3_21_18 0.00179 0.00001 0.28327 0.00001 17.1 0.9 23% D3_21_25 0.00124 0.00001 0.28325 0.00000 16.4 0.2 17% D3_21_26 0.00134 0.00001 0.28330 0.00004 18.3 2.8 18% LA-ICP-MS isotope analyses were conducted at the University of Florida. FC-1 standard yielded an external error of 1.4 ɛHf (2 standard deviation) and long-term average of -22.6 ɛHf.

a Samples highlighted in blue indicate data with high 176Hf/177Hf internal error (1 standard deviation, 0.00013-0.00024); errors are an order of magnitude greater than the rest of the observed values. See text for further discussion.

b ɛ(0) is used because the amount of ingrowth due to time (average age of 1.47 Ma; Grimes et al., 2011) only increases ɛHf by an average of 0.026 epsilon units.

c Samples highlighted in red indicate data with analytical inconsistencies and data with greater than 40% isobaric interference correction. Percent Correction = 176/177(raw)-176Hf/177Hf (corrected) /176Hf/177Hf (raw). Individual spots up to ~40% correction do not increase the range in epsilon Hf values observed. Individual spots with correction factors greater than 40% expand the limit of ranges and are excluded from averages. 68

ɛHf values today vs. ɛHf values from 0.48-2.45Ma 25.0

20.0

15.0 ) ( t

ɛ y = 1.0014x + 0.0018 10.0 R² = 0.9997

5.0

0.0 0.0 5.0 10.0 15.0 20.0 25.0 ɛ(0)

Figure 20: Comparison diagram of ɛ(0) and ɛ(t). When plotted against each other, ɛ(0) and ɛ(t) have a slope of 1, indicating that the correction factor is too small to modify ɛ(0) at the level of 1 decimal place. In ɛ(t), t is the 206Pb/238U ages for individual zircon grains collected by Grimes et al. (2011a). Ages range from 0.48 to 2.45 Ma. 69

Table 3: ɛHf range and average for MAR sample sites (0-40% correctiona) SAMPLE Low High Δ (high-low) Arithmetic 2 SD N ɛHf ɛHf ɛHf Average 1270D (all) 13.3 16.8 3.5 14.6 2.3 16 70D-19 13.3 16.8 3.5 14.8 2.7 9 70D-25 13.3 16.4 3.1 14.4 1.9 7 1275D (all) 15.4 22.1 6.7 18.4 1.9 44 75D-10 18.1 22.1 4.0 20.1 2.3 15 75D-134 16.7 17.7 1.0 17.5 0.6 9 75D-144 16.1 19.2 3.1 17.8 2.6 4 75D-180 15.4 18.8 3.4 17.3 2.2 10 75D-200 17.6 19.0 1.4 18.2 1.0 6 1309D (all) 13.3 20.7 7.4 16.6 3.6 40 1309D-58 13.3 15.6 2.3 14.3 1.4 7 1309D-250 14.6 17.4 2.8 16.0 1.8 9 1309D-564 15.2 19.3 4.1 17.0 2.8 10 1309D-820 18.1 20.7 2.6 19.1 2.2 4 1309D-1327 14.8 19.2 4.4 17.0 3.4 6 1309D-1415 14.8 19.0 4.2 17.2 3.6 4 30˚N South Wall (all) 13.7 19.1 5.4 16.1 2.8 34 3646-1205 13.7 16.7 3.0 15.1 1.7 10 3652-1333 13.8 18.2 4.4 15.6 3.6 6 3647-1359 15.7 19.1 3.4 16.9 2.1 7 D3-21 15.5 18.9 3.4 16.7 2.3 11 Atlantis Massif (all) 13.3 20.7 7.4 16.3 3.2 74 FC-1 Standard -23.9 -21.2 2.7 -22.6 1.4 55 a Percent correction calculated from isobaric interference correction. See text for further discussion.

Figure 21: Histogram of zircon ɛHf values collected in this study. ɛHf constraints shown estimated for mantle reservoirs for comparison (E-DMM, Avg. DMM, and D-DMM) are from Workman and Hart, 2005.

Chapter 4: Discussions

4.1 Intrasample Variation: Comparing Ocean Zircon to the Hf-Standard

The Lu-Hf standard used during analysis of this study (FC-1) is from a section of

gabbroic anorthosite from the Duluth Complex (Paces and Miller, 1993). The Duluth

Complex formed as a large magma chamber with a homogeneous Hf-isotopic composition (long-term 2 SD = 2.8 epsilon units; Table 3). Within-sample variation that exceeds that of FC-1 is an indication of geologic heterogeneity, possibly from magma mixing and multiple sources within the current suite of samples examined here. Only four rocks from the MAR record intrasample ranges less than 2.7 epsilon units and the rest of the samples have ranges between 2.8 – 4.4 epsilon units. The observed ranges indicate that all rocks exhibit a range of Hf-isotope values comparable to the analytical precision defined by the standard. The ranges observed between different samples, however, are larger (~200-1400 mbsf) and extend up to 7.4 epsilon units (i.e., Core 1309D). This rock- to-rock variability constrains the Hf-isotope variability of the mantle source of the small- volume intrusions that build the crust.

4.2 Geographic Variations in Zircon and MORB ɛHf

Along the MAR, the mantle is known to be chemically and isotopically variable

(Figure 22) on the scale of an entire ocean ridge (~10,000’s kms) and individual ridge

segments (~100’s kms) (Dosso et al., 1993; Andres et al., 2004). Changes in MORB

chemistry along the northern MAR suggest a geochemical gradient of enrichment from 72

north to south (Dosso et al., 1993). The basalt compositions evolve from E-MORB in the southern 14˚N region to N-MORB in the northern 16˚N region (Dosso et al., 1991, 1993), and this transition is also recorded by the ɛHf values in cores 1270D (located south of the

15˚20’ FZ) and 1275D (located north of the 15˚20’ FZ). In addition, recent evidence suggests that the mantle is also heterogeneous on smaller scales (<1-10’s km) (Harvey et al., 2006; Portner et al., 2011).

Analysis of basalts from 15˚25’ N to 16˚ N show a range of ɛHf values of 14.6 –

17 (Chauvel and Blichert-Toft, 2001; Agrainer et al., 2005). Hafnium isotope values from zircons in core 1275D range from 15.4 – 22.1, extending the trend to more depleted values. Basalts from 14˚7’ N to 14˚89’ N show a range of ɛHf values of 11.4 – 13.5

(Andres et al., 2004; Agrainer et al., 2005), which extend below those found in core

1270D (ɛHf values of 13.3 – 16.8). In general, both the basalt and zircon ɛHf values are consistent with more depleted values in the north and more enriched values in the south.

Considering the basalt end-members sampled over this distance of ~210 km, the overall range observed is approximately 6 epsilon units. In comparison, ɛHf values at core

1275D alone (< 210 meters deep) vary over a larger range (13.3 – 22.1), showing that there is a greater variation in 210 meters of core than MORB displays over ~210 km along the ridge (Figure 22).

Mantle chemical heterogeneities, such as isotopic compositions, have been suggested to be the primary factor influencing the chemical variability of mantle melts

(Niu and Hékinian, 1997; Salters and Dick, 2002). Synthesizing the isotopic and chemical nature of the discreet intrusive melts observed within the lower ocean crust 73

helps constrain possible mantle heterogeneity at small-scales during crustal accretion.

Furthermore, local scale variability (i.e., within a short vertical drill core) provides

insights into the stacking pattern of intrusions building the igneous ocean crust.

4.3 Growth of Lower Slow-Spreading Oceanic Crust

Two end-member models for lower oceanic crustal growth comprises a gabbro

glacier model (Figure 23a) and a many-sill model (Figure 23b). In the gabbro glacier

model, the lower crust is formed from a frequently replenished, shallow axial magma

chamber (Phipps-Morgan and Chen, 1992). In contrast, the many-sill model involves crystallization of gabbro throughout the lower axial crust by still intrusions (Korenaga

and Kelemen, 1998). The latter model predicts more variable compositions, and variation

in Mg # with depth in different intrusive series. Each model predicts a very different style

of cooling which has important implications for characterizing heat flow budgets in mid-

ocean ridge environments (Phipps-Morgan and Chen, 1993; Teagle et al., 2010). The

gabbro glacier cools rapidly from the top of the magma chamber near the ridge axis;

whereas the many-sill model cools slowly by conduction of heat from the lower crust.

The ɛHf values, along with published geochemical and age variability (e.g., Mg #),

provide further evidence that the gabbroic crust in this study was built by multiple

injections of magma.

4.3.1 Heterogeneous Hafnium Isotopes along the MAR.

Cores 1270D and 1275D contain ɛHf values that are distinct from one another

(Figure 21, 24a), and a student’s t-test confirms that ɛHf values from the two cores are 74

statistically significant (p > 0.01). This observation is similar to the contrast in values for

basalts erupted north and south of the 15˚20’ FZ. The extent of ɛHf variability recorded

between the two cores indicates that distinct mantle sources were sampled for the melts generated north and south of the 15˚20’ Fracture Zone

Hole 1270D contains consistently low ɛHf values (Figure 21). Rock samples

1270D-19 and 1270D-25 suggest that the mantle source being sampled during the 40 kyr of crustal construction was homogeneous at the level of analytical precision. In contrast, core 1275D contains high ɛHf values that correspond closely to the depleted end-member

of the typical depleted MORB mantle (D-DMM, ɛHf = 25.3, Workman and Hart, 2005)

(Figure 21). Furthermore, rocks from this 210-meter core yield large variations in ɛHf between samples. The shallowest rock sample (1275D-10) defines the highest ɛHf values,

(18.1 – 22.1) and is distinct from rock samples located further down-core that have single grain ɛHf values ranging from15.4 – 19.2 (Figure 24a). Zircons from 200 mbsf are statistically distinct from zircons from 134 and 180 mbsf (p > 0.05), which indicates variability within the 1275D core complex. The elevated ɛHf values seen in sample

1275D-10 indicate that the dike in the upper portion of the core had a different, more depleted mantle source than the magma constructing the lower portion of the core (from

130 – 200 mbsf).

Based on the observed ɛHf variations down-core, the lower crust here must have been built by multiple sills, rather than from a single large magma chamber during seafloor spreading. This supports the many-sill model for crustal accretion. The similarity 75

in age of 1275D-10 to other rocks in core 1275D (Grimes et al., 2011) indicates that isotopically distinct mantle sources melted over very short time-scales. This interpretation is supported by a single study on the Re-Os isotopic characteristics of mantle peridotite drilled in the 15˚20’ FZ region, which revealed significant isotopic heterogeneity over short (10’s m) intervals (Harvey et al., 2006).

Hole 1309D and surficial samples from Atlantis Massif’s southern wall contain

ɛHf values that broadly fall within a range of 13.3 – 20.7 (Δ ɛHf of 7.4) but the range in values from individual rocks is small (Δ ɛHf of 2.3 – 4.4; Table 3) and, overall, are similar to the average depleted MORB mantle (Figures 21, 24b). Rocks from the southern wall are not statistically distinct from one another. However, differences between rock samples in 1309D are indicated by Student’s t-tests and show a down-core distributions that correlates to the three lithological supergroups previously defined using whole-rock

Mg # (Blackman et al., 2006). ɛHf values increase in supergroup I, reaches its highest value in supergroup II, and then decrease in supergroup III (Figure 24b). Distinctions between these lithologic supergroups are supported by Pb/U zircon ages that indicate differences in emplacement age up to 100 – 200 kyr (Grimes et al., 2008). The variations of ɛHf values down-core, along with Mg # and age variations, suggest that these lithologic supergroups were built from small-scale intrusive melts. This is consistent with range of ɛHf values observed at 1275D.

Neodymium isotopes analyzed previously on whole rocks from 1309D using high precision TIMS (thermal ionization mass spectrometry) approach (Delacour et al., 2008).

The Nd isotope values for gabbroic rocks in 1309D are consistent with typical depleted 76

mantle values, and were all within analytical uncertainty of one another (Figure 24b). A likely reason for the greater within-sample variability found in this study is the analytical approach. Whole rock analyses by TIMS are more precise, but require an entire sample to be ground, dissolved, and analyzed. This process removes subtle differences in isotope composition resulting from a complex, multi-stage igneous history. By analyzing single crystals, a smaller snapshot in time may be resolved (i.e., a single crystallization event), whereas plutonic crust may form by the migration of melts through a crystal mush zone.

The analyses of samples obtained by the grinding of entire rocks can represent mixtures of multiple crystallization events. Another reason Nd-isotopes record less variability is because the shorter half-life (147Sm = 106 Gyr, which decays to 143Nd) reduces sensitivity compared to Hf-isotopes.

One igneous process that might explain the variation in ɛHf within a single rock

(typical within sample variations of 1.0 – 4.4 epsilon units) samples is the mixing of magmas with different initial ɛHf values during percolation through a crystal mush zone.

Melt-reaction resulting from a similar process are envisioned based on textures and mineral compositions of troctolites that formed through melt-rock reactions at the

Godzilla Megamullion oceanic core complex (located in the Philippine Sea) (Sanfilippo et al., 2013). Thus, a single rock might not form from a single parental magma, but rather crystallize over time and react with different melts with variable compositions. While the

ɛHf supports the initial interpretation by Delacour et al. (2008) that the 1309D gabbros originated from a depleted mantle source, subtle variations in the isotope composition of that source are indicated by the results presented here. 77

In summary, zircon Hf isotope values collected from ODP cores 1270D, 1275D,

and the Atlantis Massif indicate that the lower oceanic crust was built by injections of

isotopically distinct magmas, derived from a locally variable mantle. A comparison of the

rock-average 206Pb/238U zircon ages with newly reported ɛHf values suggests that there is

no temporal relationship of the measured ɛHf values (Figure 25). The ɛHf values, accretion ages, core observations of igneous contacts, and geochemical variations (e.g.,

Mg #), are consistent with a many-sill model for ocean crustal accretion (Grimes et al.,

2008; Ildefonse et al., 2006).

4.4 Investigating Mantle Source Compositions Using Zircon

Different broad magmatic environments can be distinguished using trace element ratios of U/Yb in zircon (Grimes et al., 2007). The ɛHf values in zircon from Macquarie

Island indicate an origin of enriched and depleted MORB sources, and it was recognized that select trace element ratios correlated with these differences in ɛHf (Portner et al.,

2011). Portner et al. (2011) demonstrated that plots of U/Yb versus Y (ppm) reveal differences between zircons, and suggested that such elemental ratios are applicable for distinguishing E-MORB from N-MORB. Here, we investigate whether U/Yb ratios, as well as Yb/Nb and P/Nb ratios in zircon from the MAR along with ɛHf show differences within the mantle source.

The multielement diagram for N-MORB and E-MORB (Figure 26) (Pearce and

Stern, 2006) is arranged so that the most incompatible are located on the left and all values are normalized to average MORB using values from Sun and McDonough (1989). 78

From these patterns, it is clear that ratios of highly incompatible elements (e.g., U, Nb) to

a less incompatible element (e.g., P, Yb, Y) can be used as a proxy for mantle enrichment

or depletion (Pearce and Stern, 2006). Thus, ratios like U/Yb (Figures 19a-b) provide a method for evaluating mantle source enrichment/depletion. A test of the mantle source variations in ɛHf documented in previous sections can be made by correlating these ɛHf values with trace element ratios (P/Nb, Yb/Nb, U/Yb).

The zircon data from both sites near the 15˚20’ FZ show a weak positive correlation on a plot of P/Nb versus ɛHf (R2 = 0.5; Figure 27a). Zircons from 1270D

contain low ɛHf values and the lower P/Nb ratios (< 100) are consistent with an origin

from a more enriched mantle source. Zircons from 1275D-10 have the highest ɛHf values from this study, and the larger P/Nb ratios (>100) are consistent with the interpretation that this rock sample originated from a more depleted source (Figures 27a). In contrast, zircons from the Atlantis Massif show no correlation between ɛHf and P/Nb (R2 = 0.08).

Ratios of Yb/Nb and U/Yb show a similar, though not as well-defined, trend as P/Nb

(Figures 27b, 27c). Overall, these trace element ratios are consistent in displaying differences in mantle source between the highest (1275D-10) and lowest (1270D) ɛHf values. Trace element ratios seen in the Atlantis Massif do not display differences between mantle source. The three plots in Figure 27 contain scatter for all sample locations, which can be explained by variable igneous processes in the crust, such as fractional crystallization, and variability within mantle sources, or a combination of both.

4.5 Variation in Crustal ɛHf by Mixing of Two End-Members

The range of ɛHf observed within individual rocks, and between rocks from the

same locations, can be explained either by mixing of two compositional end-member

magmas in varying proportions deeper in the mantle, or melting of a locally

heterogeneous mantle with a range of ɛHf. In the latter case, the magnitude of local scale

heterogeneity observed would be a reflection on the variation of the mantle itself>

Preservation of isotopic compositions during ascent into the crust implies accretion

according to the many-sill accretion model. However, the difference in one locality to

another in the absolute ɛHf value observed in MORB and zircon data require larger scale

variations that can be accommodated in the many-sill model.

The isotopic composition of natural basaltic magmas sampled globally fall on a

mixing line between depleted mantle (MORB source) and enriched mantle (OIB-source)

(Zindler and Hart, 1986), and is traditionally interpreted to reflect mixing of at least two end-member sources. Following the model presented by Jeffcoat (2012), I attempt to model the effect of mixing two defined end-members on the composition of zircon in the most fractionated rocks within the lower crust, using U/Yb trace element ratios and ɛHf.

Source compositions chosen for initial trace element concentrations in the model

are average MAR E-MORB and average MAR N-MORB values defined by Klein

(2003). Initial Hf isotope values are the Workman and Hart (2005) E-DMM (11.1 ɛHf) and D-DMM (25.3 ɛHf) values which define the 2σ enriched and 2σ depleted values of 80

the average DMM. E-DMM and D-DMM ɛHf values were calculated from 176Hf/177Hf values in Workman and Hart (2005) using BSE 176Hf/177Hf from Bouvier et al. (2008).

The first step of this model involves individual fractional crystallization of the depleted and enriched mantle source end members. Fractional crystallization is first incorporated because it changes the melt composition over time, and zircons form late in fractional melting sequences. The equation used to calculate the individual trace element concentrations in zircon was developed by Jeffcoat (2012) as the following:

Czrn = ((D-(D-(Kdzrn*fzrn)))/fzrn)*Cmelt*(1-F)^(D-1) (Eqn #3) where Czrn is the concentration of the element in zircon. D is the bulk distribution

coefficient of the element from Jeffcoat (2012). Kd is the partition coefficient of the

element in zircon from Fujimaki (1986) and Bea et al. (1994). Fzrn is the modal fraction

of zircon in the cumulate: 0.01 for this model. 1-F represents the fraction of liquid

remaining; F ranges from 0.01 to 0.99 and increases by 5% increments. Assumptions

made when using this model are defined by Jeffcoat (2012) and involve inferring:

compositions of the parent liquids and residuum, accuracy of mineral partition

coefficients with changing melt compositions and temperature.

After the two end-members have individually fractionated, the fractionated melts

undergo mixing. Mixing, like variable F in the previous equation, ranges from 1% to 99%

and increases by 5% increments. Mixing is calculated for specific fractionation

proportions between the depleted and enriched end-members.

Figure 28 shows possible mixing between enriched and depleted mantle sources.

The two outer mixing curves represent extreme fractionation between the enriched and 81

depleted sources, and illustrate the most extreme bounds possible for the initial parent magmas. In order for the model to incorporate the majority of data collected in this study, both the enriched and depleted melts have to experience high amounts of fractionation

(99%) while the other end-member experiences very low amounts of fractionation (1%).

The modeled end-member curves are unable to explain the composition of about one- third of the data points shown. The middle curve represents melts mixing at equal amounts of fractionation (50%) and crosses through the majority of zircon data collected in this study.

The majority of zircon compositions can be explained by the extreme case of mixing depleted mantle derived melts with the addition of 30-80% of melt from a more enriched source. However, the mixing model curves in figure 28 show that it is not possible to explain all the variability seen within this data set by mixing enriched and depleted melts that have previously been fractionated according to the model parameters.

The degree of fractionation that end member melts experience before mixing occurs is the primary influence on zircon composition (extreme upper and lower limits are required). Samples that fall below equal degrees of fractionation might be explained by mixing from melts where the depleted end member experienced greater degrees of fractionation (up to 99%) before mixing with the less fractionated enriched melts.

Samples that fall above equal degrees of fractionation might be explained by mixing from melts where the enriched end member experienced greater degrees (up to 99%) of fractionation before mixing with the less fractionated depleted melts. 82

Twenty-eight out of a total 78 analyses (36%) fall outside of the model curves and

cannot be explained by mixing of melts with even extreme degrees of fractionation.

There are several reasons to why there is a misfit of zircon data with this model. First, the

same partition coefficients (Bea et al., 1994) are used for this model. Since the partition

coefficient is a function of temperature and melt composition this value potentially differs

between EMORB and NMORB. More importantly, the initial U/Yb of model zircon

formed from the EMORB end-member (Klein, 2003) is less than the observed values in

this zircon U/Yb data set. The values from Klein (2003) are composite calculated from

EMORB and NMORB glasses on PetDB, and cannot explain the U/Yb values collected

here. The zircon data may instead indicate that there is not a single enriched end-member

value, and the data are better explained by sources that are more geochemically variable.

Thus, the mixing of two single end-members does not fully explain the variability seen in

this data set.

In contrast to the large composite intrusions (1309D and the lower 150 m of

1275D), the most enriched (1270D) and most depleted (1275D-10) end-members for ɛHf measured in the study originate in cm-scale dikes intruding ultramafic rock. While some rocks samples down core in both 1275D and 1309D show statistically significant between-sample variability, one possible explanation is that the melts intruding the larger gabbro complexes drilled in the oceanic core complexes do undergo some degree of

homogenization through mixing of enriched and depleted mantle source end members;

this variable mixing results in the large ranges of ɛHf and their heterogeneous source

signatures. Dikes at Atlantis Massif (southern wall) are statistically indistinguishable 83

from the large pluton drilled at 1309D, and this indicates that the melts in this region are an extension of the melts from 1309D.

4.6 Significance of ɛHf Variability

It has been stated in previous sections that the cores in this study have variations in epsilon units of 3.5 (for 1270D), 6.7 (for 1275D), and 7.4 (for 1309D) epsilon units.

Considering the fact that the Duluth gabbro standard FC-1 exhibits a range of less than three epsilon units, we infer that variation greater than this is the result of real geologic variation, which can only arise due to initial variation in the mantle source of magmas building the lower crust. In this section, we discuss one possible model to explain how a variation of ~6-7 epsilon units can develop in the mantle on a local scale (<10’s km’s) beneath the MAR. We start by considering two parcels of mantle, initially with an equivalent isotopic composition. The two parcels could develop different ɛHf values through alteration of the Lu/Hf ratio by preferential melt extraction from one parcel relative to another, or local refertilization), followed by the passage of time. During a second melting event, i.e., present day decompression beneath the MAR, melting of these two parcels would produce MORB with different ɛHf. To test the feasibility of this model, the relative difference in 176Lu/177Hf ratios and amount of time needed to produce a difference in ɛHf of ~6 – 7 must be evaluated.

The amount of time necessary to develop a specified difference in 176Hf/177Hf (Δ

176Hf/177Hf) can be calculated using the basic age equation for the 176Lu-176Hf radiogenic system. The basic age equation is: 84

176 177 176 177 176 177 λt Hf/ Hf = ( Hf/ Hf)initial + ( Lu/ Hf) * (e – 1). (Eqn #4)

If we consider only the difference in isotopic ratios, and assume initial 176Hf/177Hf ratios

of the two mantle parcels are equivalent, the age equation can be reduced to:

(Δ 176Hf/177Hf) = (Δ 176Lu/177Hf) * (eλt-1) (Eqn #5)

where “Δ” refers simple to the difference between the two parcels, not an absolute value.

The solution to this equation for various 176Lu/177Hf ratios is plotted on Figure 29.

Geologically reasonable Δ 176Lu/177Hf ratios were determined based on the ranges of

published Mid-Atlantic Ridge basalt glasses (compiled from Pet DB) and the values for

EMORB and NMORB in Sun and McDonough (1989). Based on these values, the

maximum difference in 176Lu/177Hf observed for Atlantic MORB Δ is 0.02. Ratios of

0.005 (ratios on the order of 10-3) define a lower end-member. For each of the Δ

176Lu/177Hf values, eλt-1was resolved using times of 1 Ma, 10 Ma, 100 Ma, 500 Ma, 1 Ga,

2 Ga, and 4 Ga, and a λ value of 1.865 * 10-11, and then a best fit curve was plotted. The

change in ɛHf (Δ ɛHf) was then solved by normalizing Δ 176Hf/177Hf to the BSE value of

0.28279. The calculated changes in Δ ɛHf with time for different Lu/Hf ratios are plotted

in Figure 29.

These model curves in Figure 29 indicate that a ɛHf difference of 6-7 epsilon

units requires ~550 Ma of time for an initial Δ 176Lu/177Hf of 0.02, and ~2 Ga for an

initial Δ 176Lu/177Hf of 0.005 (Figure 29). Since a Δ 176Lu/177Hf ratio of 0.02 defines the

maximum difference in Atlantic MORB, we can constrain that two parcels of melt would

have to have been separated for at least ~550 Ma in order to create isotopic differences of

6-7 epsilon units. The ages of ~550 Ma and ~2 Ga are consistent with Harvey et al.’s 85

(2006) interpretation that melting of two separate mantle sources has been occurring in the Mid-Atlantic Ridge since 1.5 – 2 Ga. Thus, isotopic differences within the mantle develop and are preserved over these timescales, and are not fully homogenized during subsequent convection and overturn of the mantle. These findings indicate that variable

MORB compositions can originate from the upper mantle alone, and do not require involvement of a deeper enriched source. Mixing of different mantle reservoirs within the melting triangle beneath MORs is likely to be significant as well.

4.7 Schematic Cross Section of the Mantle beneath the MAR

Zircon isotope and trace element compositions support the interpretation that the field areas in this study were constructed by discrete intrusive melts. These melts were produced from mantle sources with ɛHf values ranging from 13.3 – 22.1 epsilon units.

Within a single drill core, ɛHf values commonly varied by a narrow range of 1.0 to 4.0 epsilon units. Consideration of ɛHf values based on the size of the igneous intrusion indicates that it is possible for larger-scale intrusions to positively affect the degree of mixing between end members.

This model envisages to explain these observations from igneous material forming the lower oceanic crust is shown in Figure 30, and depicts the relationship between mantle sources and their subsequent mixing and emplacement in the lower crust.

In contrast to the gabbro glacier model where a high-level axial magma chamber constructs the lower crust, we seen that crust forms through many small-scale intrusive melts. These small-scale intrusive melts do not homogenize (due to the absence of a long- 86

lived axial melt lens) and zircons within the melt retain the distinct mixing ratios of enriched and depleted material as the melt crystallizes. Variation of Δ ɛHf ranges from ~6

– 7 epsilon units on the scale of an entire drill core, and smaller individual rock variations

(1 – 4 epsilon units) are depicted by heterogeneous pulses of melts (Figure 30).

4.8 Future Research

This study provides the first Hf isotope constraints for drill cores 1270D, 1275D, and 1309D. Although Hf constraints in MORB are becoming more common, these are the first ɛHf from lower igneous crust along the Mid-Atlantic Ridge. The data suggest that the mantle source exhibits a finite variability locally, possibly arising from varying history of melt extraction or refertilization. Further Hf isotope constraints of zircons from other MOR locations would allow these conclusions to be tested further. Examination of

MAR zircons from other drill cores would greatly improve and expand upon the database of oceanic zircon geochemistry.

The simple geochemical modeling used to estimate the amounts of mantle source mixing is not without assumptions. A more thorough model that involves various melting scenarios (as in melting in the presence of minerals like spinel and garnet) would further refine parameters involved with mantle source mixing.

Mohns- Knipovitch Ridge

Iceland

Figure 22: Published Atlantic MORB ɛHf and ɛNd values along MAR latitudes with the averaged ɛHf and ɛNd values for ODP holes 1270D and 1275D, and the Atlantis Massif. Published ɛHf and ɛNd values were calculated from 176Hf/177Hf and 143Nd/144Nd values in Chauvel and Blichert-Toft (2001); Andres et al. (2004); Agranier et al. (2005); Debaille et al. (2006); and Pet DB using BSE 176Hf/177Hf of 0.282725 [Bouvier et al., 2008] and 143Nd/144Nd of 0.512638 [Jacobsen and Wasserburg, 1980].

Figure 23: Schematic drawing of crustal accretion models, modified from Phipps- Morgan and Chen (1993), Korenaga and Kelemen (1998). A. Gabbro glacier model. B. Sheeted many-sill model. Hafnium isotope observations from this study the many-sill model for crustal accretion.

89 1275D 1270D A. Recovery Recovery 0% 100% 0% 100% 0 0 25

) 50 f s

b 75 ( m Ultrama c

h 100

t Gabbro p

e 125 Olivine Gabbro D Diabase 150 Felsic Intrusions 175 Sample Locations 200 225 10.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 ɛHf

1309D Recovery B. 0 0% 100% 0

200 200 Basalt/Diabase 400 Oxide Gabbro ) 400 Lithologic Lithologic I Supergroup f

s Gabbro, Gabbronorite b 600 600 Olivine Gabbro ( m

h Troctolite t

p 800 800 Olivine-Rich Troctolite e

D & other ultrama cs

1000 Lithologic II Supergroup 1000

1200 1200

1400 1400

1600 Lithologic III Supergroup 0.0 5.0 10.0 15.0 20.0 25.0 30.0 Epsilon Units

ODP 1270D ODP 1275D εHf, IODP 1309D εNd, IODP 1309D

Figure 24: Zircon ɛHf values collected in this study at the corresponding rock sample depths; down-hole rock types for each core are shown for reference. A. Zircon ɛHf values for 1270D and 1275D. Down-hole rock types are from Grimes et al. (2011). B. Zircon ɛHf values and ɛNd values of gabbroic rocks in IODP hole 1309D. ɛHf values are from this study; ɛNd values are from Delacour et al. (2008). Down-hole log is from Blackman et al., 2006. 90

Averaged ɛHf and U/Pb Ages 1270D-19 3.00 1270D-25 1275D-10 2.50 1275D-134 1275D-144 2.00 1275D-180

a ) 1275D-200 ( M 1.50 1309D-564 e g

A 1309D-820 1.00 1309D-1327 1309D-1415 0.50 3646-1205 3652-1333 0.00 3647-1359 10 12 14 16 18 20 22 24 D3-21 ɛHf

Figure 25: Averaged zircon ɛHf values and averaged zircon 206Pb/238U ages for each rock sample in this study. 206Pb/238U ages are from Grimes et al. (2008). Error bars are 2σ for ages and 2 SD for ɛHf. 91

Figure 26: Reference geochemical patterns highlighting the differences between trace element concentrations in EMORB and DMORB, after Pearce and Stern (2006). Normalizing values are from Sun and McDonough (1989).

B. 1000. A. 1000. Atlantis Massif: R²=0.08

15°20’ FZ: b b N /

N R²=0.5 b / 100. 100. Y P

10. 10. 10.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 10.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 ɛHf ɛHf

C. 1. ODP 1270D

ODP 1275D (10 mbsf) b Y / 0.1 ODP 1275D (130-200 mbsf) U Atlantis Massif: IODP 1309D Atlantis Massif: Southern Wall

0.01 10.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 ɛHf

Figure 27: Zircon ɛHf values plotted against select trace element ratios of A. P/Nb, R2 values within the 15˚20’FZ (R2=0.05) indicate that data from cores 1270D and 1275D are weakly correlated, and values from the Atlantis Massif (R2=0.08) indicate that data from core 1309D and the southern wall are not correlated B. Yb/Nb, and C. U/Yb. Arrows indicate trends in source enrichment as determined from Figure 26. 93

1.00

ODP 1270D ODP 1257D (10 mbsf) ODP 1275D (130-200 mbsf) Atlantis Massif: IODP 1309D Atlantis Massif: Southern Wall ) n o c r ( z i 0.10 b Y / U

0.01 8 10 12 14 16 18 20 22 24 26 28 εHafnium (zircon)

enriched melts experience greater fractionation (E 90%, D 10%) enriched and depleted melts are experiencing equal amounts of fractionation (E & D both 50%) depleted melts are experiencing greater fractionation (D 99%, E 1%) Enriched End-Member Field Depleted End-Member Field Zircons derived from mixing Zircons derived from mixing of fractionated metls. of fractionated melts.

Figure 28: Modeling results depicting possible mixing between enriched and depleted mantle melts that had been previously fractionated. The three curves represent incremental mixing between fractionated enriched and depleted melt end members. See text for further discussion. 94

. 100

1 f H

ε 0.1 Δ 176 Lu/177Hf = 0.001

Δ Black Line: Variation of 6-7 ɛHf observed in crust-forming igneous rocks sampled in core Δ176 Lu/177Hf = 0.005 U1309D and 1275D, Mid-Atlantic Ridge Δ176 Lu/177Hf = 0.02 0.01

0.001

0.0001 0 1 2 3 4 Time (Billions of years)

Figure 29: Model depicting how long two mantle reservoirs with geologically reasonable 176Lu/177Hf ratios would have to be separated in order to develop the range of ɛHf observed within core samples from a single location. Geologically feasible Δ176Lu/177Hf ratios were estimated based on measured values in Mid-Atlantic Ridge basalt glasses (downloaded from PetDB), and from Sun and McDonough’s (1989) values for EMORB and NMORB. The black line indicates a ɛHf variation of 6-7, consistent with the variation observed in cores 1309D and 1275D. 95

Depleted Mantle Reservoir Enriched Mantle Reservoir Melts Resulting from Varying Mixtures of D and E Reservoirs Basalt Ocean

Crust

Mantle

Figure 30: Schematic illustration of the mantle underlying the Mid-Atlantic Ridge, and melts accreted from the lower crustal gabbros. Melts derived from enriched (red) and depleted (blue) mantle reservoirs mix at varying degrees to produce heterogeneous melts (purple) that construct the crust through small, intrusive pulses. 96

Chapter 5: Conclusions

ODP and IODP drilling near the 15˚20’ Fracture Zone and Atlantis Massif (30˚N)

along the Mid-Atlantic Ridge (MAR), where lower crustal gabbros and mantle peridotite

are exposed directly on the seafloor, provides an opportunity to understand the processes

responsible for creating the crust covering two-thirds of the Earth’s surface. Melts

intruded into the lower crust at slow-spreading ridges provide direct insights into the

composition of the mantle at these locations, and allow characterization of the local-scale

isotopic variability recorded.

Zircon Hf isotope values collected from ODP core1275D and the Atlantis Massif

display mantle heterogeneity at local, vertical scales. The range of ɛHf values within most

individual rocks is small and similar to the range of the Hf standard, but a greater range is

observed on the scale of an entire core. Since the Hf standard FC-1 shows a ɛHf range of

2.7, these cores are comparatively heterogeneous. The variability seen at the small, vertical scale of these cores is comparable the MORB heterogeneity observed over ~200 km laterally in the 14˚ N - 16˚ N region of the MAR. Variable mixing and the resulting heterogeneity seen between rock samples down-hole supports crustal accretion through discrete, multiple intrusions of late-stage mantle melt.

Zircon U/Yb trace element ratio compositions from the 15˚20’ FZ and Atlantis

Massif show distinct differences in mantle source. When paired with ɛHf values, zircon trace element ratios (P/Nb, Yb/Nb, U/Yb) are correlated to ɛHf values; P/Nb ratios show 97

the strongest trending that clearly distinguishes the mantle source end-members of this

dataset.

Geochemical modeling suggests zircon trace element and Hf isotope values observed cannot be explained by variable mixing between previously and extremely

fractionated enriched and depleted source end members. One factor in the degree of

isotopic homogenization following mantle melting during ascent and emplacement into

the crust is the amount of melt generated. Thin dikes intruding mantle, like 1270D and

1275D-10, contain more extreme ɛHf values than rocks taken from larger composite

plutons in 1309D and the lower part of 1275D. In both examples, the larger composite

intrusions in 1275D and 1309D contain ɛHf values that plot around average DMM; these

may reflect more thorough mixing of melts from depleted and enriched mantle sources.

An estimate of how long it would take for two parcels of mantle to deviate from

one another by ~6-7 epsilon units (assuming they had the same initial ɛHf values, and

developed different 176Lu/177Hf ratios following a melt extraction or refertilization event) show that the two parcels would have had to have been separated for at least ~550 Ma,

given an initial Δ 176Lu/177Hf ratio of 0.02. This modeling shows that long time scales on

the order of ~1 – 2 Ga are needed in order to have small isotopic differences within the

mantle.

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105

Appendix A: REE patterns for oceanic zircons within individual rock samples

A. 1270D -19 10000.00

e 1000.00 t i r d

n 100.00 o C h /

t 10.00 n e

m 1.00 e l E 0.10

0.01 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

B. 1270D -25 10000.00

e 1000.00 t i r d

n 100.00 o C h /

t 10.00 n e

m 1.00 e l E 0.10

0.01 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Figure A1: REE Patterns for samples A. 1270D-19 and B. 1270D-25 in ODP core 1270D. 106

A. 1275D -10 B. 1275D -134 10000. 10000. e e t t i i r 1000. r 1000. d d n n o 100. o 100. C h C h / / 10. 10. t t n n e e 1.0 1.0 m m e e l

l 0.10 0.10 E E 0.01 0.01 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

C. 1275D -144 D. 1275D -180

10000. e e 10000. t t i i r r

1000. d d 1000. n n o o 100. 100. C h C h / / t t 10. 10. n n e 1.0 e 1.0 m m e e l 0.10 l 0.10 E E 0.01 0.01 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

E. 1275D -200

e 10000. t i r 1000. d n

o 100. C h

/ 10. t n

e 1.0 m e

l 0.10 E 0.01 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Figure A2: REE Patterns for rock samples in ODP core 1275D taken at A. 10 mbsf, B. 134 mbsf, C. 144 mbsf, D. 180 mbsf, and E. 200 mbsf. 107

A. 1309D-564 B. 1309D-820 e

10000. e 100000. t t i i r r 10000.

d 1000. d n n 1000. o 100. o 100. C h C h /

10. / t r 10. n 1.0 n e e 1.0 m m 0.10 e e 0.10 l l E E 0.01 0.01 La Ce Pr NdSmEu GdTb DyHo Er Tm YbLu La Ce Pr NdSmEu GdTb DyHo Er Tm YbLu

C. 1309D-1175 D. 1309D-1327 e 100000. e 100000. t t i i r 10000. r 10000. d d n 1000. n 1000. o o 100. 100. C h C h / / t 10. t 10. n n e 1.0 e 1.0 m 0.10 m 0.10 e e l l E 0.01 E 0.01 La Ce Pr NdSmEu GdTb DyHo Er Tm YbLu La Ce Pr NdSmEu GdTb DyHo Er TmYbLu

E. 1309D-1415 F. 3646-1205 e 100000. e 10000. t t i i r 10000. r 1000. d d n 1000. n o o 100. 100. C h C h 10. / / t 10. t n n 1.0 e 1. e m m 0.10 e 0.10 e l l E 0.01 E 0.01 La Ce Pr NdSmEu GdTb DyHo Er TmYbLu La Ce Pr NdSmEu GdTb DyHo Er TmYbLu

G. 3652-1333 H. 3647-1359

e 100000. e 100000. t t i i

r 10000. r 10000. d d

n 1000. n 1000. o o 100. 100. C h C h / /

t 10. t 10. n n

e 1.0 e 1.0

m 0.10 m 0.10 e e l l

E 0.01 E 0.01 La Ce Pr NdSmEu GdTb DyHo Er TmYbLu La Ce Pr NdSmEu GdTb DyHo Er TmYbLu

I. D3-21 10000. e t i

r 1000. d

n 100. o

C h 10. / t

n 1.0 e 0.10 m e l

E 0.01 La Ce Pr NdSmEu GdTb DyHo Er TmYbLu Figure A3: REE patterns for rock samples from the Atlantis Massif. A. – E. are samples taken from IODP core 1309D. F. – I. are surficial samples taken from the southern wall. 108

Appendix B: Histograms, probability density functions, error-weighted mean plots, and mean squared weighted deviates for all samples collected in this study A. 26 10 error bars plotted as 2 SD 9 8 22 7 6 18

5 f H 4 ɛ 3 14 2 1 10 0 7 9 11 13 15 17 19 21 23 25 27 29 εHf 6 ODP 1270D Weighted Mean: 16.1±0.7 [4.2%] 95% conf. MSWD: 64, probability: 0.00

B. 26

4 22

3 18 f H 2 ɛ 14

1 10

0 7 9 11 13 15 17 19 21 23 25 27 29 εHf 6 1270D -19 Weighted Mean: 16.2±0.9 [5.6%] 95% conf. MSWD: 97, probability: 0.00

C. 26

6 22 5

4 18 f H 3 ɛ 14 2

1 10

0 7 9 11 13 15 17 19 21 23 25 27 29 εHf 6 1270D -25 Weighted Mean: 14.5±0.6 [3.8%] 95% conf. MSWD: 2.7, probability: 0.01

Figure B1: Single grain ɛHf results plotted as error weighted mean plots (right) and as histogram with probability density functions (left) for A. All single grain analyses from ODP Core 1270D. Analyses are organized by increasing value. The same for each sample analyzed from this core (B. and C.). Errors for ɛHf are shown as 2 SD internal errors in all plots. 109

A. error bars plotted as 2 SD 26 15 ODP Hole 1275D 22

10 18 Weighted Mean: 18.6±0.3 [1.5%] 95% conf. f

H MSWD: 57, probability: 0.00 ɛ 5 14

0 10 7 9 1113 15 1719 21 23 25 2729 εHf B. 1275D-10 6 22 5 Weighted Mean: 19.6±0.4 [1.8%] 95% conf. 4 18 f

3 H MSWD: 18, probability: 0.00 ɛ 2 14 1 0 10 7 9 1113 15 1719 21 23 25 2729 εHf C. 9 8 22 1275D-134 7 1275-1D34 6 18 Weighted Mean: 17.6±0.1 [0.5%] 95% conf. 5 f H MSWD: 0.43, probability: 0.90 4 ɛ 3 14 2 1 0 10 7 9 1113 15 1719 21 23 25 2729 εHf D. 3 1275D-144 22

2 18 Weighted Mean: 16.2±0.7 [4.3%] 95% conf. f

H MSWD: 3.3, probability: 0.02 ɛ 1 14

10 0 7 9 1113 15 1719 21 23 25 2729 εHf

E. 1275D-180 22 7 6 Weighted Mean: 18.7±0.4 [2.3%] 95% conf. 18 5 f MSWD: 64, probability: 0.00 H

4 ɛ 3 14 2 1 10 0 7 9 1113 15 1719 21 23 25 2729 εHf F. 1275D-200 4 22 3 Weighted Mean: 18.2±0.6 [3.0%] 95% conf. MSWD: 0.31, probability: 0.91

f 18

2 H ɛ 14 1

0 10 7 9 1113 15 1719 21 23 25 2729 εHf

Figure B2: Single grain ɛHf results plotted as error weighted mean plots (right) and as histograms with probability density functions (left) for A. All single grain analyses from ODP Core 1275D. Analyses are organized by increasing value. Zircons analyzed from 1275D are shown as blue bars- note that in core 1275D are derived from this one sample (B.-F.). Analyses grouped by host rock. Errors for ɛHf are shown as 2 SD internal errors in all plots. 110

error bars plotted as 2 SD A. IDOP hole 1309D 10 22 8 18 Weighted Mean: 16.5±0.3 [2.0%] 95% conf. 6 f

H MSWD: 41, probability: 0.00 ɛ 4 14 2 0 10 7 11 15 19 23 27 εHf B. 4 1309D-58 22 3 18 Weighted Mean: 14.1±0.5 [3.9%] 95% conf. f

H MSWD: 1.7, probability: 0.12 2 ɛ 14 1 10 0 7 11 15 19 23 27 εHf C. 1309D-250 5 22 4 18 Weighted Mean: 16.5±0.7 [4.1%] 95% conf. f

3 H

ɛ MSWD: 118, probability: 0.00 2 14 1 0 10 7 11 15 19 23 27 εHf

D. 1309D-564 22 5 4 18 Weighted Mean: 16.4±0.8 [4.9%] 95% conf. f

3 H ɛ MSWD: 20, probability: 0.00 2 14 1 10 07 11 15 19 23 27 εHf E. 22 1309D-820

4 18 f Weighted Mean: 19.1±1.7 [9.0%] 95% conf.

3 H

ɛ MSWD: 7.0, probability: 0.00 2 14 1 0 10 7 11 15 19 23 27 εHf F. 22 1309D-1327 2 18 f Weighted Mean: 17.5±1.2 [7.1%] 95% conf. H ɛ MSWD: 8.9, probability: 0.00 1 14

10 07 11 15 19 23 27 εHf G. 22 1309D-1415 2 18 f Weighted Mean: 16.8±2.3 [14%] 95% conf. H ɛ MSWD: 1.4, probability: 0.23 1 14

10 0 7 11 15 19 23 27 εHf Figure B3: Single grain ɛHf results plotted as error weighted mean plots (right) and as histograms with probability density functions (left) for A. all single grain analyses from IODP Core 1309D (analyses are organized by increasing value) and each sample analyzed from this core (B.-G.). Errors for ɛHf are shown as 2 SD internal errors in all plots. 111

A. error bars plotted as 2 SD 25 22 Atlantis Massif 20 15 18 f Weighted Mean: 15.5±0.2 [1.3%] 95% conf. H 10 ɛ MSWD: 62, probability: 0.00 14 5 0 10 7 11 15 19 23 27 εHf

B. 12 Atlantis Massif, Southern Wall 22 10 8 18 Weighted Mean: 15.1±0.2 [1.1%] 95% conf. f

6 H MSWD: 26, probability: 0.00 ɛ 4 14 2 10 0 7 11 15 19 23 27 εHf C. 5 22 3646-1205 4 18 3 f Weighted Mean: 15.2±0.3 [2.1%] 95% conf. H

ɛ MSWD: 3.0, probability: 0.00 2 14 1 10 0 7 11 15 19 23 27 εHf D. 22 3652-1333 2 18 Weighted Mean: 14.6±0.8 [5.2%] 95% conf. f H

ɛ MSWD: 4.6, probability: 0.00 1 14

10 0 7 11 15 19 23 27 εHf E. 5 22 3647-1359 4 18 Weighted Mean: 16.6±0.6 [3.4%] 95% conf. 3 f H

ɛ MSWD: 3.1, probability: 0.01 2 14 1 10 0 7 11 15 19 23 27 εHf F. 22 D3-21 7 6 18 5 f Weighted Mean: 15.1±0.3 [1.9%] 95% conf. H

4 ɛ MSWD: 65, probability: 0.00 3 14 2 1 0 10 7 11 15 19 23 27 εHf

Figure B4: Single grain ɛHf results plotted as error weighted mean plots (right) and as histograms with probability density functions (left) for all samples collected at A. all single grain analyses from the Atlantis Massif, B. all single grain analyses from the surficial southern wall, and each sample analyzed from the southern wall (C.-F.). Errors for ɛHf are shown as 2 SD internal errors in all plots. ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !

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