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IDENTIFYING MAGMA RECHARGE

USING CRYPTIC MINERAL VARIATIONS

IN THE SOMERSET DAM IGNEOUS

COMPLEX, QUEENSLAND, AUSTRALIA

Brenainn Simpson

Bachelor of Science

Submitted in fulfilment of the requirements for the degree of Master of Applied Science

School of Earth, Environmental and Biological Sciences

Science and Engineering Faculty

Queensland University of Technology

2017

i

Keywords

Electron microprobe, geochemistry, igneous petrology, in situ mineral chemistry, layered

mafic intrusion, layered gabbro, laser ablation inductively coupled plasma mass

spectrometry, Somerset Dam Igneous Complex.

ii

Abstract

The Somerset Dam Igneous Complex (SDIC) layered gabbro is one of the few moderately exposed

mafic intrusive rocks in Queensland, approximately 3km across and 1km thick. Previous research on

the SDIC interprets the layered gabbro to be comprised of seven macro-layers (called macrocycles).

This thesis presents data for 20 major and trace elements from plagioclase, olivine and pyroxene

collected via electron microprobe and laser ablation inductively coupled mass spectrometry. This

chemical stratigraphy allows for the investigation of major and trace element distribution in the

layered gabbro. A large troctolite unit between 245m and 325m above sea level contains plagioclase

characterised by common reverse zoning, a marked change in An component, anomalous trace

element distributions and localised flow alignment. These features are interpreted to be the result of

magma recharge disrupting the magmatic evolution of the complex. The existing model for the SDIC

interprets each macrocycle as a magma recharge event within a single chamber. As the signs of

magma recharge occur within one of the proposed macrocycles, rather than at the contact between the

macrocycles, it is proposed that the SDIC layered gabbro is the product of several stacked sills

forming a composite sill. The SDIC has broad implications for identifying the occurrence and effect

of magmatic recharge in mafic systems.

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Table of Contents

Keywords .................................................................................................................................. i

Abstract .................................................................................................................................... ii

Table of Contents .................................................................................................................... iii

List of Figures ......................................................................................................................... iv

List of Tables ............................................................................................................................ v

Statement of Original Authorship ........................................................................................... vi

Acknowledgements ................................................................................................................ vii

1. Introduction ............................................................................................................. 2

Study aims ................................................................................................................................. 4

2. The Somerset Dam Igneous Complex ................................................................... 6

The layered series ..................................................................................................................... 7

3. Experimental Methods ......................................................................................... 10

Electron microscopy ............................................................................................................... 12

Laser ablation inductively coupled mass spectrometry .......................................................... 14

4. Results .................................................................................................................... 15

Petrography and phase chemistry ........................................................................................... 15

Rock Descriptions ................................................................................................................... 18

Leucogabbro ................................................................................................................. 21

Troctolite ...................................................................................................................... 24

Olivine gabbro .............................................................................................................. 29

Oxide gabbro ................................................................................................................ 31

Minor and trace element chemistry ......................................................................................... 33

5. Discussion............................................................................................................... 40

Thermal re-equilibration ......................................................................................................... 42

Controls on the chemical profiles of plagioclase and olivine ................................................. 43

Variations in phase chemistry and rock texture ...................................................................... 44

What constitutes a macrocycle? .............................................................................................. 50

Composition of ‘recharge’ magma ......................................................................................... 53

SDIC – Revised model ............................................................................................................ 57

6. Conclusion ............................................................................................................. 63

References .................................................................................................................. 65

iv

List of Figures

Figure 2.1 Geological map of the SDIC pg.9

Figure 2.2 Composite stratigraphic diagram of the SDIC layered gabbro pg.9

Figure 3.1 Map of sampling locations pg.11

Figure 4.1 Example plagioclase demonstrating classification scheme pg.17

Figure 4.2 Lithofacies textures in the SDIC pg.18

Figure 4.3 Stratigraphic diagram for sampling area pg.19

Figure 4.4 Plutonic rock classification diagram after the International Union of

Geological Sciences pg.20

Figure 4.5 Pyroxene quadrilateral classification diagram pg.23

Figure 4.6 Troctolite rock textures pg.26

Figure 4.7 Plagioclase reabsorption textures pg.27

Figure 4.8 Troctolite core to rim chemistry plots pg.28

Figure 4.9 Eu content in SDIC plagioclase pg.33

Figure 4.10 Ba vs Eu content in SDIC plagioclase pg.34

Figure 4.11 Ba vs An component for major rock types pg.34

Figure 4.12 Minor element core to rim plots for the SDIC pg.36

Figure 4.13 Minor element core to rim plots for the central troctolite unit pg.36

Figure 4.14 Rare earth element spidergrams for the SDIC pg.39

Figure 5.1 Model for the evolution of the SDIC after Mathison (1967) pg.42

Figure 5.2 Diffusion length scales for major mineral phases pg.44

Figure 5.3 Core normalized rare earth element spidergrams for the central troctolite unit pg.49

Figure 5.4 Variation in An component with height modified after Mathison (1987) pg.52

Figure 5.5 Rare earth element spidergrams for modelled liquid compositions pg.55

Figure 5.6 Revised model for the evolution of the SDIC pg.61

v

List of Tables

Table 3.1 Sampling locations pg.10

Table 3.2 List of standards for electron microprobe pg.14

Table 4.1 Plagioclase classification scheme for the SDIC pg.15

Table 4.2 Classification of all sampled plagioclase pg.16

Table 4.3 Moderal mineralogy of thin sectioned samples pg.20

Table 4.4 Leucogabbro plagioclase mineral chemistry pg.21

Table 4.5 Variation in olivine compositions in the SDIC pg.22

Table 4.6 Olivine mineral chemistry in the SDIC pg.22

Table 4.7 Troctolite plagioclase mineral chemistry pg.25

Table 4.8 Olivine gabbro plagioclase mineral chemistry pg.30

Table 4.9 Oxide gabbro plagioclase mineral chemistry pg.32

Table 4.10 Trace element analyses for SDIC plagioclase pg.37

Table 5.1 Calculated partition coefficients for modelled liquid compositions pg.53

Table 5.2 Modelled trace element compositions for liquids in equilibrium with sampled

plagioclase cores and rims pg.56

vi

Statement of Original Ownership

I, Brenainn Simpson, affirm that the work contained in this thesis has not been previously submitted to meet requirements for an award at this or any other higher education institution. To the best of my knowledge and belief, the thesis contains no material previously published or written by another person except where due reference is made.

Signature

QUT Verified Signature

Brenainn Simpson

Date: November 2017

vii

Acknowledgements

I’d like to acknowledge and sincerely thank my supervisors, Professor David Gust and Dr Patrick

Hayman. Additionally I’d like to thank Dr. Charlotte Allen, Miss Karine Harumi Moromizato and Dr.

Henrietta Cathey at CARF for their assistance and training, Mr. Michael McFadyen for granting

access to his land, Gus Luthje for making thin sections, and finally Mr. Mathew Beutell, Mrs. Janelle

Simpson and Mr. Marko Uksanovic for assisting with field work and sample collection. Thank you to

all who have supported me in this undertaking from start to finish.

2

1. INTRODUCTION

The earliest ideas underpinning modern understanding of magmatic systems can be traced back to the

research pioneered by Bowen (1928) and subsequent authors (eg. Fenner, 1929; Wager and Deer,

1939), much of which utilised layered mafic instrusions (LMIs) as the subject. LMIs are intrusive

igneous rocks of mafic or ultramafic chemistry that have characteristic and often repetitive layering

from which they derive their name. LMIs have long been considered the best case studies of

magmatic differentiation due to their characteristic sequence of repeating layers, which are easily

observed and mapped on the macro level. These intrusions vary in size, age and chemical

composition. The Skaergaard intrusion in East Greenland and the Bushveld in South Africa can be

considered end member examples of layered mafic plutonism. The Skaergaard intrusion, a relatively

small (~100km2) mafic intrusion emplaced 55 million years ago (Nielsen 2004) contrasts to the very

large (approximately 66,000km2) and ancient (2.06Ga) Bushveld igneous complex (Wager and Brown

1968). The smaller scale and closed system of the Skaergaard make it an ideal body for studying

basaltic plutonism (e.g., McBirney and Hunter, 1995; McBirney, 1998; Nielsen, 2004; Tegner, 1997).

In his seminal paper ‘The Evolution of Igneous Rocks’, Bowen (1928) coined the term ‘fractional

crystallisation’ to explain a simple evolution of magma as elements are locked into a solid crystalline

structure. Over time the magma cools, forming solid phases and enriching the residual liquid in

elements which are incompatible with minerals under the ambient pressure and temperature regime.

The evolution of LMIs may also involve magmatic replenishment with hot new mafic or ultramafic

liquids mixing and mingling with the existing evolved liquid to a variable degree. It is the addition

and mixing of these new liquids that have yet to evolve by fractional crystallisation that provides a

convenient explanation for the formation of layers in these bodies (Brown 1956; McBirney and Noyes

1979; Irvine 1980; Turner 1980). As a magma evolves via fractional crystallisation, it cools and

enriches in elements that are not used in the formation of minerals. Over time this results in a cooler

magma of a new or ‘evolved’ composition which directly alters the species and composition of the

growing minerals. However, magma recharge resets the fractionation sequence of minerals to a less

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evolved stage and multiple injections of new melt can be responsible for new sequences of layers that

then evolve via fractional crystallisation. The exact nature of this process is fertile ground for

continuing research in both igneous petrology and fluid dynamics (e.g., (Marsh 2013; Latypov et al.

2015).

The stratigraphy of many LMIs and in particular, of the Skaergaard intrusion, demonstrates that

fractional crystallisation should generate layers that are more enriched in Si, Na and Fe with height

(Wager and Deer 1939). Building upon Bowen’s early work regarding mineral segregation (Bowen

1915), Wager and Brown (1968) argued that gravity was the primary driver separating minerals based

on their density to explain the layering observed in the Skaergaard intrusion. This model proved

insufficient to explain many observations. For instance, the accumulation of dense minerals such as

olivine at the roof of the intrusion with lower density minerals such as plagioclase being beneath it

was anomalous. Revisions of the model of gravity settling included consideration of the governing

role of fluid dynamics within the chamber. McBirney and Noyes (1979) demonstrate that as a mafic

melt cools and crystallises it quickly becomes non-Newtonian in character and considered double-

diffusive convection to be a dominant mechanism of the formation of the layering. The double-

diffusive convection model is strongly supported by laboratory experiments and data obtained in the

field (Turner 1980; Rice 1981; Huppert and Turner 1981; Wyborn et al. 2001). The effects of

compacting the crystalline mush (thus ejecting evolved interstitial fluids to re-mix or mingle with the

main chamber are addressed by (Irvine 1980). (DePaolo 1981) models the impact of country rock

assimilation on evolving melts with implications as to how these processes can disturb a predicted

liquid line of descent. The nature of in situ crystallisation versus mineral transport itself is considered

by (Langmuir 1989) to be of critical importance for identifying the nature of physical processes in the

magma chamber. Thus a model that considers not only fractional crystallisation but also fluid

dynamics and compaction is able to better test hypotheses regarding the evolution of an intrusive

igneous rock.

An additional layer of complexity arises when considering the interactions between multiple

magmatic liquids in a semi-open or open system. The behaviour of major and trace elements in an

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open system can be very different from that of a closed system depending on the magnitude and

timescale of exposure to new magmas. (O’Hara 1977) presents research on an extreme open system

end-member – a mid-ocean ridge. Continuous recharge with a similar magma produces a steady state

system preventing whole scale differentiation by a liquid line of descent as predicted by fractional

crystallisation (O’Hara and Mathews 1981; O’Hara and Fry 1996). Mixing of two chemically distinct

magmas, typically mafic and felsic melts can result in a mixture of a new intermediate composition

(Kouchi and Sunagawa 1985). Experimental studies (e.g., Zimanowski et al., 2004) demonstrate the

immiscible nature of this mixing resulting in an incomplete mingling of the liquids.

The partitioning of minor and trace elements into minerals (Cawthorn 2007; Toplis et al. 2008;

Tanner et al. 2014) aids in identifying the effects of magma recharge in the chamber. Modern

analytical techniques allow for the collection of chemical data on the micron scale from the core to the

rim, documenting the crystallisation of the magma chamber. Like much of the seminal work in

igneous petrology, these studies focus on LMIs for documenting the impacts of magma recharge in

the evolution of igneous rocks.

Study Aims

This study aims to identify magma recharge by using mineral composition to better understand this

process in the formation of mafic layered rocks. Similar studies have been published on large layered

intrusions (Cawthorn 2007; Tanner et al. 2014) but this work is complicated by subsolidus re-

equilibration, which obscures the magmatic signatures. By focusing this research on the Somerset

Dam Igneous Complex (here after SDIC), an intrusion of considerably less volume and less complex

thermal history, this study attempts to avoid the obstacles resulting from thermal re-equilibration.

Previous research on the SDIC focussed on field mapping, petrology, and geochemistry describing the

macro-scale structure of the intrusion and offering petrogenetic explanations (McLeod 1956;

5

Mathison 1970; Walsh 1972; Riley 1991; Walker 1998). Mathison and Walker propose fractional

crystallisation and gravity settling disturbed by incidents of magma recharge as the mechanisms

behind the formation of the layers. This study tests this hypothesis by using mineral major and trace

element chemistry correlated with stratigraphic height to identify chemical disequilibrium attributable

to magmatic replenishment. This work is possible due to to the significant advances made in recent

years in analytical equipment and procedure. In particular, EMPA allows the construction of high-

resolution mineral core to rim chemical profiles (Cawthorn 2007; Tanner et al. 2014) and LA-ICPMS

enables a slightly lower resolution profile of trace element variation.

Plagioclase is targeted in this study as chemical variations are readily inferred from petrographic

examination. The substitution of Ca and Na- cations, which is coupled with the substitution of Al and

Si cations, is a kinematically slow process which can result in chemically distinct zones in the crystal.

Various zonation patterns (normal, reversed, oscillatory) along with other textural evidence (sieve

textures, ragged cores, etc.) can be attributed to a variety of different processes (changes in pressure,

water content, temperature and liquid composition). By ruling out some ‘environmental’ parameters,

these patterns may provide a clear signal of changes in liquid composition (‘magma recharge’).

Finally, plagioclase is a commonly occuring cumulate phase in mafic layered intrusions which

provides the best potential for identifying chemical evidence of magmatic processes.

6

2. THE SOMERSET DAM IGNEOUS COMPLEX

The SDIC is a Late Triassic (225 ± 4 Ma based on 40Ar/39Ar and 87Rb/87Sr (Walker (Walker 1998))

intrusion located west of Somerset in south-east Queensland, approximately 115 kilometres north-

west of Brisbane (Figure 2.1). It is a part of a belt of Late Triassic plutons distributed across the

southern section of the northern New England Orogen (Purdy 2013a). Described as a high alumina

tholeiitic gabbro (Riley (1991), the SDIC is one of the few mafic intrusions in Queensland and is

generally obscured by cover with some moderately well-exposed areas. The SDIC is a shallow

intrusion (<5km) (Mathison, 1987) that is crystallised from magmas already highly fractionated in a

deeper chamber (Riley, 1991) based on low ppm concentrations of Ni and Cr. It intrudes the middle

Triassic Neara Volcanics, which comprise a series of intermediate volcanic material within the

Toogoolawah Group (Purdy 2013b) of the Esk Trough.

The Esk Trough developed in the Permian-Triassic in three stages (Hill 1960; Campbell 2005). The

basin was formed as a result of Early Permian extension followed by a period of thermal subsidence

and finally foreland loading during the latest Permian-Early Triassic. Deposition in the basin

comprised sequence of older marine sediment overlain by the terrestrial Toogoolawah Group which

includes the Neara Volcanics. Exposed highlands associated with the Hunter-Bowen orogeny

provides the provenance for the sediments of the Toogoolawah Group and indicates that orogeny had

produced an uplift in the region before the Mid Triassic (Campbell 2005). The tectonic history of the

Esk Trough correlates chronologically, lithologically, stratigraphically and in paleoenvironment with

the eastern Bowen Basin. Regional scale contraction associated with the Hunter-Bowen orogeny is

considered to have resulted in the westward migration of arc magmatism onto the continent resulting

in the emplacement of volcanoclastic material and Late Permian-Mid Triassic I-type intrusions

(Campbell 2005)

At the broadest scale, the SDIC is comprised of a mafic layered series of approximately 500 meters

(Mathison, 1987) to 600 meters (Walsh 1972; Riley 1991) in thickness with a chilled margin at the

7

sides (mapped as a ‘fine-grained gabbro’ by Walker (1998). It is capped by diorite and granite

(McLeod 1959a; Mathison 1967). The earliest research on the SDIC (McLeod, 1959a) did not

consider the overlying intermediate and felsic rocks as being related to the layered gabbro. Walker

(1998) demonstrates that these rocks are fractionated products of the gabbro based on εNd values. The

mafic layered series is approximately circular in plan view and 3km in diameter with planar features

in the layering that dip gently towards the centre of the intrusion, approximately 10-15o. The SDIC

has not undergone significant metamorphism or deformation (McLeod 1959a; Mathison 1970; Riley

1991; Walker 1998) and has not been overturned. This is supported by the internal structuring of the

layers which are more evolved at the top (in the oxide gabbro) and more primitive at the base (in the

troctolite).

While the SDIC is moderately well exposed, many areas are entirely under cover or otherwise

inaccessible. The base and roof of the layered series used in this study are obscured by cover and the

overlying diorite.

The Layered Series

The mafic layered series of the SDIC is divided into seven macrocycles with each cycle starting with

an leucogabbro, followed by a troctolite, an olivine gabbro and an oxide gabbro (McLeod 1959a;

Mathison 1967; Riley 1991; Walker 1998). The third macrocycle is an exception to this sequence and

is considered to be an ‘interrupted’ macrocycle by Mathison (1967, 1987). The subdivision of the

macrocycles is based on modal mineralogy and bulk rock composition. Cryptic variations within each

macrocycle are expressed by changes in plagioclase (An%, K2O % and Sr ppm) and olivine (Fe2O3 %

and Mn ppm) compositions (Figure 2.2). In total there are 22 distinct units defined by sharp phase,

modal and textual contacts (Figure 2.2). Macrocycles 1, 2, 3 and 4 are approximately 80 – 100m in

thickness while the uppermost cycles (5, 6, and 7) are much thinner (20 – 30m).

The earliest research published on the SDIC considers the complex as a whole to be a typical ‘gabbro

– red rock association’ (McLeod, 1959; Grout, 1918). McLeod (1959) states that the ‘red’ orthoclase-

8

rich granite and quartz diorite are not differentiation products of the gabbro and were emplaced

‘slightly later’ than the gabbro. The mafic layered series or ‘banded gabbro’ is interpreted to be a

product of gravity settling with each macrocycle beginning with a troctolite and progressing to an

olivine gabbro, followed by an oxide gabbro and finally a leucogabbro (Mathison, (Mathison 1967).

The oxide gabbro is more correctly classified as a Fe-Ti oxide gabbro, however the term oxide gabbro

is used in this thesis for consistency with the literature. Mathison (Mathison 1987) revised the

sequence of rock types in each macrocycle, to begin with, the leucogabbro. This revision is based on

the observation that while the troctolite is typically the most primitive unit, the leucogabbro contains

common reversely zoned plagioclase and represents the interaction between new and old liquids

resulting in the chemical resetting of the liquid in the crystallisation front to a more primitive

chemical composition. The reverse zonation is interpreted as evidence to support a model where each

macro-cycle is the product of a new injection of primitive magma combined with fractional

crystallisation and gravity settling.

In a typical sequence the leucogabbro is followed by the troctolite (+olivine), then the olivine gabbro

(+pyroxene) and finally the oxide gabbro (-olivine +magnetite). Walker (Walker 1998) describes the

complex as comprising a fine-grained gabbro followed by a succession of rhythmically layered

gabbro, overlain in turn with intermediate composition rock facies and finally a felsic granophyre.

Mathison (1987) observes that each macrocycle has an individual chemical evolution rather than a

simple progression from most primitive at the base to most evolved at the roof. This is attributed to

the variable chemical character of the liquid in the crystallisation front which is dependent on the

degree and composition of the individual recharge pulses. Each macrocycle begins with a sharp

contact and is proposed to represent the injection of a new pulse of melt entering a single magma

chamber, overlying the existing solid components of the chamber and crystallising into a new layer.

9

Figure 2.1. Map of the Somerset Dam Igneous Complex after (Walker 1998). Macrocycles are labelled 1 – 7. Sampling

traverse for this study is indicated beginning in the second macrocycle, through the third and ending in the fourth.

Figure 2.2. Stratigraphic column showing the inferred cyclic units taken from (Mathison 1987). Sampling area for this study

is indicated to the right with a dotted line. Height given is above base which is estimated to be approximately 20-30m below

seas level. Macrocycle boundaries are indicated with a dotted line and the cycle units number 1-7.

10

3. EXPERIMENTAL METHODS

Thirty-seven samples were collected in a traverse across ~250m of stratigraphy with absolute height

measurements taken via GPS coordinates (Table 3.1, Figures 2.1 & 3.1). This traverse provides

almost continuous exposure of the mafic layered series. It begins in the oxide gabbro unit of

macrocycle 2, continues across macrocycle 3 and finishes in the oxide gabbro layer of macrocycle 4

(Figure 2.2). Twenty-one thin sections were made for a petrographic study being cut perpendicular to

the dominant direction of the layering if present. High-resolution maps of each thin section were

created using a Leica DM6000. Data collection took place at the Central Analytical Research Facility

(CARF) located at the Queensland University of Technology Garden’s Point campus.

Table 3.1. Latitude, longitude, and height for collected samples. Height is measured relative to sea level (MASL) (Australian

Height Datum). Height relative to base (MAB) is estimated to be ~25m below sea level after Mathison (1987).

Sample Name Locations (GDA 94) Classification

Revised # Latitude Longitude MASL (m) MAB (m) Rock name

BSSD 36 -27.115 152.53542 414 389 oxide gabbro


BSSD 32 -27.11636 152.53508 403 378 olivine gabbro


BSSD 28 -27.11761 152.535 382 357 troctolite

BSSD 27 -27.11783 152.53489 377 352 leucogabbro




BSSD 19 -27.1205 152.53592 325 300 troctolite

BSSD 16 -27.12128 152.53556 304 279 troctolite

BSSD 15 -27.12139 152.53547 296 271 troctolite

BSSD 13 -27.122 152.53472 276 251 troctolite

BSSD 11 -27.12236 152.53456 257 232 troctolite




BSSD 06 -27.12358 152.53369 188 163 troctolite


11

Figure 3.1. Topographic map of the research area (GDA 94 MGA Zone 56). Contours are at 20m intervals. Samples are

labelled 1 – 36. The sampled traverse shown here is placed into the greater context of the SDIC in Figure 2.1.

12

Electron microscopy

Preliminary SEM analysis of polished thin sections was conducted on a Hitachi TM3000 equipped

with an energy-dispersive X-ray spectroscopy (EDS) with an accelerating voltage of 15 kV. Data

collected using the TM3000 informed quantitative analyses collected via EPMA and is not published

here. High-resolution backscatter electron (BSE) images were acquired using the JOEL JXA-8530 F

electron probe micro analyser (EPMA).

EPMA data was collected using a JOEL JXA-8530 F electron microprobe with a take-off angle of 40

degrees, an accelerating voltage of 15 kV and a probe current of 20-25 nA. Elements were acquired

using analysing crystals LIFH Kα for Ti, Cr, Mn, Fe, Ni, PETL Kα for Ca, TAP Kα for Al, Si, and

TAPH Kα for Na, and Mg.

The locations of the analysing spots were placed carefully to avoid any form of mineral alteration.

Mineral set ups and standards

Plagioclase analyses were conducted with a 10 µm beam diameter, and pyroxene and olivine analyses

were collected with a 3 µm beam diameter. The counting time for all mineral set ups was 10 seconds

on peak and 10 seconds off peak for measuring Kα for Na, 20 seconds on peak and 20 seconds off-

peak measuring Kα for Si, Mg, Al, K, 25 seconds on peak and 25 seconds off-peak measuring Kα for

Ca, 30 seconds on peak and 30 seconds off-peak measuring Kα for Mn, Ti, Cr, Ni and Fe and 40

seconds on peak and 40 seconds off peak for Ba, Rb, and Sr. The off-peak correction method was

linear for all elements.

A complete list of standards used for EPMA collection can be found in Table 3.2.

Detection limits and corrections

Detection limits were in the order of 0.2 weight percent for major element while analytical sensitivity

(at the 99% confidence level) ranged from 0.1 to 0.5 percent. Oxygen was calculated by cation

stoichiometry and included in the matrix correction. The matrix correction method was ZAF or Phi-

13

Rho-Z utilising the Armstrong/Love Scott algorithm (Armstrong 1988). Data reduction was achieved

using the Probe for EPMA software (Donovan et al. 2016).

Table 3.2. Astimex Microanalysis Standards and National Bureau of Standards (NBS) Glass standards used in this study.

Standard Element

Albite Astimex block position 1 Na (Kα)

Barite Astimex block position 5 Ba (lα), S (Kα)

Celestite Astimex block position 13 S (Kα), Sr (lα)

Diopside Astimex block position 21 Ca, Mg, Si (Kα)

Hematite Astimex block position 25 Fe (Kα)

Na-Andesite Glass Standards Block position 22 Na (Kα)

NBS K411 Glass Standards Block position 11 Mg (Kα)

NBS K412 Glass Standards Block position 10 Ca, Mg, Si (Kα)

Orthoclase Astimex block position 41 Al, K (Kα)

Oxide synthetic Astimex block position 17 Cr (Kα)

Pentlandite Astimex block position 36 Ni (Kα)

Plagioclase An59 Astimex block position 35 Al, Ca, Si (Kα)

Plagioclase An65 Astimex block position 35 Al, Si (Kα)

RbTiPO5 Astimex block position 52 Rb, Ti (Kα)

Rhodonite Astimex block position 39 Mn (Kα)

Rutile Astimex block position 40 Ti (Kα)

Willemite Astimex block position 46 Zn (Kα)

14

Laser ablation inductively coupled plasma mass spectrometry

A suite of rare earth elements (La, Ce, Pr, Nd, Sm, Eu, Gd, Dy), as well as Sr, Ba, and Rb, was

collected from in situ plagioclase grains from selected samples from the traverse of the mafic layered

series of the SDIC. ICP-MS data was collected with an Agilent 8800 with laser ablation. The

instrument uses two quadrupoles with an intervening reaction cell and a single collector. For these

analyses, the quadrupoles were synchronised with no reaction gas. An Ar gas flow at 1 litre per

minute was used with a plasma power of 1350 Watts. Laser ablation was conducted using an ESI New

Wave ArF laser with a wavelength of 193 nm. A 45µm spot was collected for each analysis with a

fluence of 2 mJ/cm2 and a repetition rate of 5-10Hz. The ablation takes place in a He atmosphere at

550 ml/min. Each analysis consists of 30 seconds for background measurement with the laser off

followed by 30 seconds of data acquisition.

Detection limits for large ion lithophile elements varied from 0.01ppm to a maximum of 0.2ppm for

Sr and Ba and 0.4ppm for Rb and had a typical relative error of approximately 15% for Rb and 5% for

Sr and Ba. Detection limits for the lanthanides ranged from 0.01ppm to 0.03ppm for La, Pr, Sm, Eu,

Dy and up to 0.06ppm for Ce, Nd and Gd. Lanthanides heavier than Gd were below detection limits.

Relative errors ranged from 4.5% to 93% for La, Ce, Pr, Sm, Nd, Eu with the average error being

approximately 10%. Relative errors for Gd and Dy are significantly greater ranging from 8.75% to

200% averaging approximately 35%.

Using high-resolution BSE images, the locations of the EPMA analyses were marked and targeted for

analysis by LA-ICP-MS. The quantitative EPMA data was used as an internal standard for reducing

the ICP-MS analyses (Ca in plagioclase). Samples were also standardised using the National Institute

of Standards and Technology Standard Reference glasses 610 and 612 at regular time intervals no

greater than 20 minutes. The data reduction is achieved using Iolite published by the University of

Melbourne. A more detailed discussion of the techniques used in the acquisition of this LA-ICP-MS

data can be found in (Longerich et al. 1996).

15

4. RESULTS

Petrography and phase chemistry

The rocks that comprise the SDIC are dominantly composed of plagioclase and olivine with minor

pyroxene. The rocks vary from leucogabbro, troctolite, olivine gabbro and oxide gabbro. Plagioclase

is the most abundant mineral being present in all rocks in abundances of 50% or greater. Detailed

microscopic examination of the grains coupled with extensive EPMA analysis allows plagioclase to

be grouped into different types (Table 4.1). This grouping is used in the petrographic description of

the rocks.

Plagioclase is grouped by zonation pattern. Plagioclase is either normally-zoned (Type 1), reversely-

zoned (Type 2), or zoned in an oscillatory manner. Normally and reversely-zoned plagioclase that is

also oscillatory zoned have been grouped separately (Types 3 and 4). Representative textures and

zonation patterns of the different plagioclase groups illustrate the combinations of Type and Group

(Figure 4.1). Classification for SDIC plagioclase are found in Table 4.2.

Table 4.1 Plagioclase are grouped based on the following criteria.

Classification Description Type 1 Normally-zoned plagioclase with low amounts of variationType 2 Reversely-zoned plagioclase with low amounts of variationType 3 Oscillatory zoned plagioclase that is overall normal in character Type 4 Oscillatory zoned plagioclase that is overall reverse in character Type 5 Unzoned plagioclase Group A Concentric zonation patternGroup B Patchy or irregular areas of contrasting compositionsGroup C Isolated pockets of chemically distinct materialGroup D No observable zonation

For example, plagioclase 06-1 from sample SD06 (Figure 4.1a) is a normally-zoned plagioclase (Type

1) that is not concentrically zoned but instead has zonation that is patchy or uneven (Group B). In

contrast, plagioclase 15-1 from sample SD15 (Figure 4.1b) is reversely-zoned (Type 2) with a

concentric zonation pattern (Group A). Plagioclase 27-5 from sample SD27 and 02-7 from sample

SD02 (Figures 4.1c and d respectively) show some oscillation in their core to rim compsition.

However, 27-5 is overall normally-zoned while 02-7 is overall reversely-zoned. These are examples

16

of Type 3 and Type 4 plagioclase. Finally, plagioclase 36-3 from sample SD36 (Figure 4.1e) is an

unzoned plagioclase (Type 5) with an isolated patch of chemically distinct material (Group C, shown

in the grey scale as a dark patch close to the rim of the grain). Group D plagioclase have either no

zonation or no zonation that is observed under BSE or polarised light.

The term ‘patchy zoning’ describes zonation patterns are irregularly distributed in contrast to

concentric zoning (Vance 1962, 1965; Streck 2008). Terms such as ‘normal’ or ‘reverse’ zoning

cannot be directly applied to plagioclase with patchy zoning since the relationships between growth

phases of the mineral can be difficult to establish. With this in mind, the core to rim analyses of many

patchy zoned plagioclase have been collected and characterised as either ‘over all normal’ or ‘over all

reverse’ in character. In this thesis BSE images of plagioclase have been included, where possible, to

provide the necessary context for zonation in the plagioclase. Samples that contain a significant

proportion of concentrically zoned plagioclase are rare in the SDIC, with the exception of SD 15

(Table 4.2). A full suite of BSE images collected for plagioclase is available in the electronic

appendix.

Table 4.2. Plagioclase classification for all samples. A blank entry indicates that this sample was not used for data collection. Sample SD16 is missing classification by zonation pattern due to this sample not being BSE imaged.

1 2 3 4 4b 5 5b 6 7 8 9 10 11

SD36 2 C 4 B 5 C 1 C 5 C 5 C 5 C

SD35 4 B 5 D 4 D

SD32 1 B 2 B 3 B

SD30 5 D 3 A 3 A 3 A

SD28 3 A 4 A 4 A 4 B 3 A

SD27 3 B 1 B 3 B 4 B 3 B 3 B 1

SD25 3 B 3 C 3 C 1 B

SD22 4 B 4 A 4 B 3 B

SD21 1 B 1 D 3 A 3 C 3 A 1

SD19 1 D 3 B 3 B 3 A 1 B 3 B

SD16b 3 B 1 B 1 B 1 B 1 B 3 B

SD16 3 3 1 3 2 4

SD15 2 A 4 A 4 A 4 A 4 A

SD13 4 A 3 B 4 B 2 B 4 A

SD11 4 A 3 A 4 B 4 B 3 B 3 B 2 B

SD07 3 B 3 A 3 B 3 A 4 B

SD06 1 B 3 B 4 B 4 B 1 A 4 B 3 B

SD02 4 B 2 B 4 A 4 B 4 B 3 B

17

Figure 4.1. a – e from top to bottom: A. Type 1B (SD06 plag-1) B. Type 2A (SD15- plag-l1) C. Type 3B (SD27 plag-5) D. Type 4A (SD02 plag-7) E. Type 5C (SD36 plag-3). Core values are plotted at the origin of the X-axis (0 µm) and progress towards the rim as distance increases. On the BSE images, cores are marked with a C and rims marked with an R. Greyscale reflects composition with lighter shades indicating areas richer in Ca and darker shades indicating areas richer in Na.

18

Rock Descriptions

For the purpose of describing and classifying the rock types of the SDIC the original cumulate

classifications (ortho, meso, ad after Wager et al. 1960) is combined with descriptions of the

cumulates by their dominant mineralogy as described by Irvine (1982). Representative

photomicrographs of the four rock types that comprise the layered series (Figure 4.2) show that the

mineralogy and textures vary from leucogabbro adcumulates to pyroxene and olivine-rich

mesocumulates. A stratigraphic column of the studied sequence (Figure 4.3) summarises rock type

and associated plagioclase and olivine mineral composition. Poikilitic textures are present in all

observed rock types. Plagioclase and olivine are the dominant chadacrysts and pyroxene, hornblende

and magnetite are commonly oikocrysts in the cumulate. The samples have undergone limited

alteration with only minor chlorite and talc alteration present as well as minor orthopyroxene reaction

rims associated with some olivine. Additionally, plagioclase typically contain varying amounts of

sericite alteration in all sampled rock types. The samples are classified as leucogabbro, troctolite,

olivine gabbro and oxide gabbro (Table 4.3, Figure 4.4). The oxide gabbro is a gabbro with a large

modal proportion of cumulus phase oxide which is typically magnetite.

Figure 4.2. (A) Leucogabbro adcumulate with minor brown hornblende; (B) Troctolite cumulate under cross-polarised light

with a strong lineated texture; (C) Olivine gabbro with augite and minor hornblende. (D) Oxide gabbro characterised by

cumulus phase magnetite with plagioclase and hornblende. Mineral abbreviations after (Siivola and Schmid 2007).

19

Figure 4.3. Stratigraphic column for sampled stratigraphy of the SDLS. The sampling area is indicated in Figure 2.2 with

boundaries between macrocycles 2,3,4 indicated. Range of plagioclase variation (An%) is shown for core (black) and rims

(red), and for olivine (Fo%) variation (black). The median is indicated with a dot. Key: Upper oxide gabbro (UOxG), Upper

olivine gabbro (UOlG), Upper troctolite (UT), Upper leucogabbro (UL), Central troctolite (CT), Lower oxide gabbro

(LOxG), Lower Troctolite A (LTa), Lower leucogabbro A (LLa), Lower Troctolite B (LTb), Lower Leucogabbro B (LLb).

20

Table 4.3. Modal mineralogy of samples collected via point counting (n=300).

Rock Type Sample Plagioclase Olivine Clinopyroxene Oxides Hornblende

# % # % # % # % # %

Oxide Gabbro SD 36 148 49.33 5 1.67 90 30 45 15 12 4

Oxide Gabbro SD 35 127 42.33 16 5.33 108 36 48 16 1 0.33

Olivine Gabbro SD 32 211 70.33 37 12.33 39 13 4 1.33 9 3

Olivine Gabbro SD 30 197 65.67 15 5 46 15.33 5 1.67 37 12.33

Troctolite SD 28 211 70.33 65 21.67 23 7.67 1 0.33 0 0

Leucogabbro SD 27 243 81 6 2 17 5.67 6 2 28 9.33

Leucogabbro SD 25 262 87.33 5 1.67 2 0.67 2 0.67 29 9.67

Leucogabbro SD 22 247 82.33 6 2 1 0.33 10 3.33 36 12

Leucogabbro SD 21 261 87 7 2.33 12 4 3 1 17 5.67

Troctolite SD 19 224 74.67 52 17.33 13 4.33 6 2 5 1.67

Troctolite SD 16 228 76 33 11 37 12.33 2 0.67 0 0

Troctolite SD 15 237 79 44 14.67 17 5.67 2 0.67 0 0

Troctolite SD 13 206 68.67 70 23.33 10 3.33 2 0.67 12 4

Troctolite SD 11 234 78 51 17 15 5 0 0 0 0

Leucogabbro SD 07 272 90.67 0 0 0 0 8 2.67 20 6.67

Troctolite SD 06 217 72.33 55 18.33 1 0.33 18 6 9 3

Oxide Gabbro SD 02 224 74.67 0 0 47 15.67 29 9.67 0 0

Figure 4.4. Classification of mafic intrusive rocks after LeMaitre and International Union of Geological Sciences, 2005.

21

Leucogabbro (SD 02,07,21,22,25,27)

The leucogabbros are adcumulates with 80% to 90% plagioclase, <5% olivine and variable amounts

of hornblende, pyroxene and magnetite. Some chlorite alteration is present in the leucogabbros, along

with the sericite alteration that is present in all sampled rock types. Plagioclase and hornblende are

poikilitic commonly enclosing oxide phases. Plagioclase oikocrysts average 2.5mm in length with

some being up to 4mm. Plagioclase is commonly normally-zoned (Type 1 and 3) from An 40 to An

65 (Table 4.4). These zonation patterns are patchy or uneven. However, some concentric patterns are

present. SD22 contains several samples that are reversely-zoned.

Olivine is euhedral to subhedral and measures approximately 0.25mm in length. It is compositionally

homogeneous (Fo69-68; Table 4.5 and 4.6) and comprises less than 5 percent of the mode. Pyroxene

is an interstitial phase in the cumulate, comprising much less than 5% of the sample or being absent

altogether. Pyroxene varies from Ca44.5 Mg43.5 Fe12.0 to Ca43.5 Mg45 Fe11.5 (Figure 4.5).

Table 4.4. Selected major element (wt%) core and rim analyses for leucogabbro plagioclase. Samples are representative of

the general character of the lithofacies. The full data set can be found electronically in Appendix 1.

07-1 07-2 21-1 22-1 22-5 27-1

Oxide Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim

SiO2 54.25 58.48 55.71 56.41 51.72 54.91 54.16 53.54 52.22 53.83 51.67 53.41

Al2O3 28.37 26.12 27.74 27.52 28.77 27.9 28.8 29.33 29.81 28.91 30.05 29.15

FeO 0.32 0.24 0.29 0.3 0.5 0.35 0.42 0.32 0.45 0.28 0.4 0.4

CaO 10.94 8.15 10.15 9.73 11.93 10.73 11.5 12.01 13.07 11.72 13.28 12.04

Na2O 4.94 6.47 5.25 5.44 3.66 5.02 4.4 4.24 3.7 4.44 3.5 4.34

K2O 0.31 0.43 0.29 0.27 0.17 0.23 0.33 0.23 0.22 0.23 0.15 0.24

Total 99.4 100.1 99.6 99.8 98.4 99.3 99.8 99.8 99.7 99.5 99.3 99.8

Si 2.47 2.62 2.52 2.54 2.38 2.49 2.45 2.43 2.38 2.44 2.37 2.43

Al 1.52 1.38 1.48 1.46 1.56 1.49 1.54 1.57 1.6 1.55 1.62 1.56

Fe 0.01 0.01 0.01 0.01 0.02 0.01 0.02 0.01 0.02 0.01 0.02 0.02

Ca 0.53 0.39 0.49 0.47 0.59 0.52 0.56 0.58 0.64 0.57 0.65 0.59

Na 0.44 0.56 0.46 0.47 0.33 0.44 0.39 0.37 0.33 0.39 0.31 0.38

K 0.02 0.02 0.02 0.02 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01

O 8 8 8 8 8 8 8 8 8 8 8 8

Total 12.99 12.98 12.97 12.97 12.89 12.98 12.97 12.98 12.98 12.98 12.97 12.98

Ab 44.15 57.49 47.52 49.48 35.31 45.23 40.13 38.45 33.4 40.11 31.99 38.93

An 54.03 40.01 50.74 48.93 63.59 53.38 57.92 60.18 65.27 58.51 67.1 59.67

Or 1.82 2.5 1.74 1.59 1.1 1.39 1.96 1.36 1.33 1.38 0.91 1.4

22

Table 4.5. Fo component variation in olivine in the SDIC. Olivine ranges in compositions from Mg1.3Fe0.7SiO4 (Fo65) to

Mg1.5Fe0.5SiO4 (Fo75).

Olv1 Olv2 Olv3 Olv4

Rock-type Sample Min Max Min Max Min Max Min Max

Oxide Gabbro SD36 67.58 68.36 68.52 69.63 67.68 67.68 - -

Olivine Gabbro SD32 70.5 70.72 70.77 71.76 72.62 72.62 71.11 71.38

Olivine Gabbro SD30 69.48 70.16 - - - - - -

Troctolite SD28 71.89 72.01 72.28 72.41 72.35 72.49 72.05 72.27

Leucogabbro SD27 69.29 69.29 69.7 70.05 69.49 69.62 69.8 69.8

Leucogabbro SD25 68.9 69.05 - - 68.87 68.87 - -

Leucogabbro SD22 69.99 70.21 69.28 69.89 69.78 69.97 - -

Leucogabbro SD21 71.63 71.63 72.66 72.66 - - - -

Troctolite SD19 74.21 74.31 73.99 74.16 73.91 74 73.89 74.12

Troctolite SD16 72.48 72.71 71.38 71.53 72.44 73.01 72.95 73.2

Troctolite SD16b 74.28 74.59 73.49 73.68 - - - -

Troctolite SD15 74.23 74.5 73.51 73.52 - - - -

Troctolite SD13 74.68 75.82 - - - - - -

Troctolite SD11 73.65 73.95 - - 73.32 73.62 - -

Troctolite SD06 66.53 66.77 66.59 66.88 66.39 66.54 66.08 66.13

Table 4.6. Selected olivine mineral composition for the SDIC reported as wt% oxide and calculated stoichiometry. Sample

progress in stratigraphic height from left to right.

06-1 15-1 16-3 27-2 30-1 36-1

Oxide Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim

SiO2 37.68 37.41 38.17 38.12 37.64 37.42 36.86 36.9 36.73 36.82 36.92 37.19

TiO2 - - 0.04 0.03 0.08 0.02 0.02 0.06 0.02 0.01 - -

FeO 29.02 29.11 23.12 23.5 24.99 24.99 27.26 27.24 27.45 27.05 29.15 28.53

MgO 32.43 32.46 37.88 38.04 37.12 36.81 35.5 35.28 35.45 35.47 34.09 34.42

MnO 0.53 0.58 0.4 0.43 0.47 0.46 0.54 0.54 0.53 0.52 0.52 0.51

CaO 0.07 0.04 0.07 0.06 0.06 0.04 0.06 0.04 0.07 0.04 0.05 0.08

Total 99.8 99.6 99.8 100.2 100.4 99.7 100.2 100.1 100.3 99.9 100.7 100.7

Si 1.01 1.01 1 1 0.99 0.99 0.98 0.99 0.98 0.98 0.99 0.99

Fe 0.65 0.66 0.51 0.51 0.55 0.55 0.61 0.61 0.61 0.6 0.65 0.64

Mg 1.3 1.31 1.48 1.48 1.46 1.45 1.41 1.4 1.41 1.41 1.36 1.37

Mn 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

O 4 4 4 4 4 4 4 4 4 4 4 4

Total 6.98 6.99 7 7 7.01 7.01 7.02 7.01 7.02 7.02 7.01 7.01

23

Figure 4.5 Pyroxene quadrilateral. The bulk of the data has been highlighted using an insert box for clarity. Troctolite

samples are highlighted with a red circle, leucogabbros in blue, oxide gabbro in green and olivine gabbro in purple.

24

Troctolite (SD 06,08,11,13,15,16,19,28)

Troctolites are adcumulate to mesocumulates and comprise 70 to 75% cumulus plagioclase and 10%

to 15% olivine with <7.5% of interstitial pyroxene and ~1% accessory magnetite. The olivine are

typically fresh, although some have experienced significant talc alteration. The troctolite facies

exhibits varying degrees of preferred alignment and is the only facies sampled that has this texture

(Figure 4.6). Samples SD11 through SD19 comprise the ‘central troctolite unit’ (CT, Figure 4.3) that

is approximately 80m thick. The central troctolite unit is characterised by a change in grain size and

texture when comparing the top and bottom samples (SD11 and SD19) with the middle samples

(SD13, SD15, SD16). At the base, the plagioclase grains are larger (up to 3mm) and are not orientated

along a preferred axis (Figure 4.6) while samples SD15 and SD16 contain smaller laths on average (1-

2mm). SD15 and SD16 also have a strong preferred orientation. SD19 is similar to SD11 with an

increase in the size of individual plagioclase and no apparent alignment of the grains.

Plagioclase from the central troctolite commonly include cores of other plagioclase that have been

absorbed into the new mineral (Figure 4.7). When a low An plagioclase comes into contact with a

more primitive liquid that plagioclase will crystallise a rim that is more An rich than the core and

become reversely zoned. Magma that is more primitive can also be much hotter and result in some

remelting of the transported plagioclase crystal. This results in a plagioclase with a core that has a

sharp change in An composition that has an anhedral form characterised by embayments and an

amorphous geometry as shown in Figure 4.7.

In addition to textural variation, the central troctolite unit also has a change in plagioclase zonation

pattern with stratigraphic height (Figure 4.8). The plagioclase in the lower section of this unit (SD11-

15) is reverse zoned, often within an oscillatory nature and a concentric zonation pattern (Type 4A)

(Table 4.2). The plagioclase in these units are less An-rich at the core (An43-76, average core value

An60). The plagioclase in the upper section of the unit (SD16-19) is normally-zoned and has uneven

patterns (Type 1B and 3B). These plagioclase are more An-rich than the plagioclase from the lower

section (An 50-80, average core value An70). Major element chemistry for troctolite plagioclase is

summarised in Table 4.7.

25

Olivine are Fo rich (Fo72-75. Table 4.5), equidimensional and subhedral to rounded. Within a sample,

the size of the olivine is consistent. However, there is a significant variation in size between the

crystals in the stratigraphically lower troctolites which can be as small as 0.25mm and the

stratigraphically higher troctolites which contain crystals up to 4mm. Where the olivine grains are

small, they are also more numerous such that the modal component of olivine within all troctolites are

similar. Pyroxene is present as diopsidic augite and varies from Ca46.9 Mg41.9 Fe11.2 to Ca44.8 Mg44.4

Fe10.7 (Figure 4.5). Where present the poikilitic pyroxene forms a mesocumulate with olivine and

plagioclase. Other areas of the cumulate where pyroxene is absent is characterised by adcumulate

plagioclase broken up by the occasional olivine.

While SD 16 and 28 contain slightly more pyroxene than a true troctolite they are very similar in

nature to the other samples in this facies.

Table 4.7. Major element (wt%) analyses for troctolite plagioclase core and rims. Samples 11 – 19 are taken from the central

troctolite unit. Samples 06 and 28 are taken from the lower and upper troctolite units.

11-1 13-1 15-3 16b-4 19-1 06-1 28-2

Oxide Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim

SiO2 52.15 51.20 52.41 51.45 54.39 50.70 48.02 49.63 52.27 44.57 53.07 55.03 52.10 52.07

Al2O3 29.77 30.84 29.98 30.73 28.39 31.03 32.51 31.12 29.47 25.18 29.49 28.79 30.59 30.62

FeO 0.45 0.44 0.43 0.42 0.31 0.46 0.44 0.28 0.41 6.04 0.32 0.34 0.48 0.42

CaO 12.63 13.56 12.61 13.34 10.97 13.97 15.96 11.00 12.62 7.23 11.95 10.83 13.10 13.31

Na2O 3.98 3.44 3.98 3.48 4.76 3.14 2.11 3.37 3.83 2.85 4.68 5.40 3.73 3.70

K2O 0.17 0.08 0.18 0.12 0.16 0.16 0.07 1.72 0.17 0.07 0.20 0.15 0.16 0.14

Total 99.4 99.7 99.8 99.7 99.1 99.6 99.3 97.3 99.0 94.2 99.9 100.7 100.3 100.5

Ions

Si 2.38 2.34 2.38 2.34 2.47 2.32 2.22 2.32 2.40 2.41 2.41 2.47 2.36 2.35

Al 1.60 1.66 1.61 1.65 1.52 1.67 1.77 1.72 1.59 1.57 1.58 1.52 1.63 1.63

Fe 0.02 0.02 0.02 0.02 0.01 0.02 0.02 0.01 0.02 0.02 0.01 0.01 0.02 0.01

Ca 0.62 0.66 0.61 0.65 0.53 0.68 0.79 0.55 0.62 0.68 0.58 0.52 0.64 0.64

Na 0.35 0.30 0.35 0.31 0.42 0.28 0.19 0.31 0.34 0.27 0.41 0.47 0.33 0.32

K 0.01 0.00 0.01 0.01 0.01 0.01 0.00 0.10 0.01 0.00 0.01 0.01 0.01 0.01

O 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00

Total 12.99 12.98 12.98 12.98 12.97 12.98 12.99 13.01 12.97 12.96 13.00 13.00 12.98 12.98

Ab 35.91 31.28 35.93 31.86 43.57 28.65 19.19 31.85 35.07 41.37 40.99 47.00 33.66 33.21

An 63.05 68.21 62.98 67.40 55.44 70.37 80.36 57.45 63.89 57.97 57.87 52.14 65.37 65.97

Or 1.03 0.50 1.08 0.74 0.99 0.98 0.44 10.70 1.04 0.65 1.14 0.86 0.98 0.83

26

Figure 4.6. Photomicrographs of the central troctolite unit (SD11 – SD19). Samples SD 13, 15 and 16 are finer-grained and

are strongly flow aligned.

27

Figure 4.7. BSE imaged plagioclase from SD13 (within the central troctolite) that features a reabsorbed Na rich core,

surrounded by a new Ca rich growth and finally an Na rich rim. The greyscale image is lighter where the plagioclase is more

Ca-rich and darker when the plagioclase is more Na-rich.

28

Figure 4.8. a – d. Selected troctolite plagioclase variation diagrams with the corresponding backscatter electron images.

Samples progress in stratigraphic height moving bottom (SD13) followed by SD15 and SD16 then SD 19 (top). SD13 and

SD15 commonly contain reversely zoned plagioclase while SD16 and SD19 are typically normally-zoned and more enriched

in An component. Sampling traverses are marked with white crosses.

29

Olivine Gabbro (SD 30,32)

Olivine gabbros are cumulates composed of 65 to 70% plagioclase, 5 to 10% olivine, and 15%

clinopyroxene as well as variable amounts of interstitial brown hornblende. Sericite alteration is

common along with some chlorite. Where hornblende is present, it is generally associated with the

alteration of pyroxene. Plagioclase range in size from 1mm – 3mm and is either normally-zoned or are

unzoned. Core compositions vary from An60 – 65. These plagioclase are either weakly concentrically

zoned or have no observable zonation. Major element chemistry is summarised in Table 4.8

Olivine is euhedral, measuring up to 3mm, and comprise 5 - 10% modal abundance. Compositionally

they are more evolved than the olivine sampled from the troctolite but less evolved than olivine

sampled from the leucogabbros (Fo70 – 72, Table 4.4). Pyroxenes occur as either poikilitic augite or

enstatite reaction rims associated with olivine. Clinopyroxene (augite) vary in composition from Ca44.5

Mg42.7 Fe12.8 to Ca43.0 Mg45.3 Fe11.7 (Figure 4.5). Pyroxene is often enclosed by hornblende while also

enclosing plagioclase and olivine itself. The dominant lithofacies of the olivine gabbro is an olivine-

plagioclase-pyroxene-hornblende mesocumulate. However, there are areas where plagioclase forms

an adcumulate texture.

30

Table 4.8. Major element (wt%) analyses from plagioclase from the upper olivine gabbro unit (UOlG, Figure 4.3) (SD30 –

32). Samples are representative of the lithofacies.

30-1 30-2 32-1 32-2 32-3

Oxide Core Rim Core Rim Core Rim Core Rim Core Rim

SiO2 53.46 53.40 53.24 55.19 53.08 54.24 56.44 53.98 52.70 55.94

Al2O3 29.99 29.87 29.97 28.66 29.08 28.25 27.00 28.56 29.38 27.98

FeO 0.39 0.38 0.42 0.30 0.44 0.38 0.33 0.43 0.43 0.50

CaO 12.39 12.34 12.43 11.04 12.20 11.17 9.63 11.49 12.41 10.26

Na2O 4.17 4.15 4.08 4.82 4.15 4.71 5.49 4.63 3.98 5.31

K2O 0.23 0.22 0.24 0.29 0.27 0.28 0.48 0.23 0.27 0.27

Total 100.9 100.5 100.6 100.6 99.5 99.2 99.5 99.6 99.5 100.5

Si 2.40 2.41 2.40 2.48 2.42 2.47 2.55 2.45 2.41 2.51

Al 1.59 1.59 1.59 1.52 1.56 1.52 1.44 1.53 1.58 1.48

Fe 0.01 0.01 0.02 0.01 0.02 0.01 0.01 0.02 0.02 0.02

Ca 0.60 0.60 0.60 0.53 0.60 0.54 0.47 0.56 0.61 0.49

Na 0.36 0.36 0.36 0.42 0.37 0.42 0.48 0.41 0.35 0.46

K 0.01 0.01 0.01 0.02 0.02 0.02 0.03 0.01 0.02 0.02

O 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00

Total 12.98 12.98 12.98 12.97 12.98 12.98 12.98 12.98 12.98 12.98

Ab 37.34 37.34 36.73 43.36 37.47 42.60 49.33 41.59 36.14 47.59

An 61.33 61.34 61.87 54.91 60.93 55.76 47.84 57.02 62.22 50.85

Or 1.33 1.32 1.40 1.72 1.60 1.64 2.83 1.39 1.64 1.56

31

Oxide Gabbro (SD 10,35,36)

The oxide gabbro lithofacies (Mathison 1987) is a mesocumulate composed of 50% plagioclase, 30%

to 35% pyroxene and 15% oxide phase(s). Plagioclase commonly contain sericite alteration, and the

augites are associated with brown and green hornblende alteration. Plagioclase is less abundant than

in the other units and is typically subhedral. It is approximately 3mm in size with a range from 0.5mm

– 5mm. The composition of the plagioclase varies from An55 to An60 with outliers as low as An40.

Variation within each grain is limited with many crystals exhibiting weak or no zonation. Major

element chemistry for oxide gabbro plagioclase is found in Table 4.9.

The other dominant mineral phase in the oxide gabbro is augite. The augites are a subhedral cumulate

phase which varies in size from 1 to 3mm. They are characterised by a higher Fs component and vary

from Ca43.5 Mg41.8 Fe14.6 to Ca39.2 Mg45.0 Fe15.8 (Figure 4.5). Olivine is rare in these samples, typically

0.5mm in size comprises 2-3% of the total rock volume and more Fe-rich (Fo68-69, Table4.4).

Oxide phases can be either early stage euhedral magnetite that can be Cr or Ti rich or later stage

intercumulate anhedral magnetite. Euhedral magnetite is typically 0.1mm – 0.5mm in size and occurs

commonly as oikocrysts in olivine and plagioclase. Primary magnetite phases are commonly included

inside pyroxene chadacrysts and are euhedral and 0.25mm in size. Late stage oxide phases comprise

the vast majority of interstitial cumulus material along with minor brown hornblende.

32

Table 4.9. Selected major element (wt%) analyses for oxide gabbro plagioclase cores and rims. Samples selected are

representative of the character of their lithofacies.

02-1 25-3 36-1 36-2 36-3

Oxide Core Rim Core Rim Core Rim Core Rim Core Rim

SiO2 55.96 53.24 53.55 53.48 54.10 54.13 54.10 54.77 54.15 54.61

Al2O3 27.71 29.21 29.16 29.33 28.81 29.02 28.92 29.10 28.74 28.62

FeO 0.33 0.39 0.39 0.35 0.39 0.35 0.40 0.39 0.34 0.33

CaO 9.88 11.87 11.97 11.95 11.45 11.67 11.54 11.53 11.35 11.33

Na2O 5.60 4.54 4.36 4.35 4.79 4.46 4.62 4.60 4.64 4.82

K2O 0.42 0.28 0.20 0.18 0.23 0.22 0.22 0.13 0.23 0.14

Total 100.1 99.8 99.8 99.8 100.1 100.0 100.0 100.7 99.7 100.1

Si 2.52 2.42 2.43 2.43 2.45 2.45 2.45 2.46 2.45 2.47

Al 1.47 1.56 1.56 1.57 1.54 1.55 1.54 1.54 1.54 1.52

Fe 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01

Ca 0.48 0.58 0.58 0.58 0.56 0.57 0.56 0.55 0.55 0.55

Na 0.49 0.40 0.38 0.38 0.42 0.39 0.41 0.40 0.41 0.42

K 0.02 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

O 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.00

Total 12.99 12.99 12.98 12.98 12.99 12.98 12.98 12.97 12.98 12.98

Ab 49.38 40.26 39.23 39.26 42.50 40.37 41.49 41.63 41.93 43.11

An 48.20 58.13 59.57 59.64 56.19 58.34 57.22 57.61 56.73 56.04

Or 2.41 1.61 1.20 1.10 1.31 1.29 1.29 0.76 1.34 0.85

33

Minor and trace element compositions

The macro scale distribution of trace elements in the SDIC varies based on position within the

stratigraphy. Plagioclase samples from below the large central troctolite unit contain different trace

element compositions from samples obtained from within the central troctolite and above. Analyses

for Eu content demonstrate the distinct chemical difference when plotted against An component

(Figure 4.9). The Eu content of samples within and above the central troctolite falls between 0.5 and

1.0 ppm and do not show significant variation as An content changes. The majority of Eu analyses

collected from below the central troctolite fall between 1 and 2ppm. There is little overlap between

the two populations. Similarly, as Ba content increases in the samples Eu content also increases

(Figure 4.10). The samples that were taken from below the central troctolite show a larger increase in

Eu content relative to Ba than the samples taken from within and above the central troctolite.

Although Sr is a compatible element in plagioclase it does not show similar chemical behaviour to Eu

in the sampled plagioclase.

Figure 4.9. Eu content in plagioclase (ppm) vs. mol % An content. Eu analyses from within and above the central troctolite

plot between the 0.5 and 1.0 lines (dotted). Eu analyses from below the central troctolite plot between 1 and 2ppm.

34

Figure 4.10. Ba content in plagioclase (ppm) vs. Eu content (ppm). Eu analyses from within and above the central troctolite

plot to the left of the dotted line and Eu analyses from below the central troctolite plot to the left.

Figure 4.11. Ba content (ppm) vs. mol % An component. Samples have been selected that represent typical trace element

distributions.

35

Fe, Ba and Sr zoning in plagioclase

Minor element zoning patterns in SDIC plagioclase are often different to their major element

variations. Selected trace element data quantifying the minor and trace element variations have been

compiled (Table 4.11). Representative core to rim plots (Figure 4.12) for Fe, Ba and Sr presenting

zonation patterns of each of the four main plagioclase groupings used to categorise the main element

data (Types 1, 2, 3 and 4) allow for several key observations to be made.

Plagioclase from the SDIC that are not sampled from within the central troctolite are typically Type 1

and 2 plagioclase with less amounts of Type 3 and 4 plagioclase and have limited or gradual variation

in their An component zonation patterns (Figure 4.12a and b). In these samples Ba has a general

pattern of decreased concentration when An component increases. Both Fe and Sr typically increase

concentration with An component.

In contrast to the other samples, samples comprising the central troctolite unit (SD 11, 13, 15, 16, 19)

have variations in minor element zonation patterns that are often complex. In unzoned or normally

zoned plagioclase (Figure 4.13c and d) Ba has a less obvious relationship with changes in An content

from core to rim. Similarly, Fe and Sr maintain relatively constant concentrations of these elements

and have little or no variation as An component changes. Reversely zoned plagioclase have similar

anomalous trends to the normally zoned plagioclase. Fe content still increases with changes in An

component although large changes in An component correspond with only minor increases in Fe

(Figure 4.13a,b,d).

36

Figure 4.12a – d (left to right). Fe, Ba and Sr vs. An bivariate plots of representative plagioclase for main plagioclase types.

Type 1 (left) through Type 4 (right). Analytical spots are shown to scale (10 µm for EPMA and 45 µm for LA-ICP-MS).

Distances from core to the rim are plotted on the X axis with the core set at the origin.

Figure 4.13a – d (left to right). Fe, Ba and Sr vs. An bivariate plots of selected samples from SD 15, SD 16, and SD 19.

Analytical spots are shown to scale (10 µm for EPMA and 45 µm for LA-ICP-MS). Distances from core to the rim are

plotted on the X axis with the core set at the origin.

37

Table 4.10. Selected trace element chemistry for SDIC plagioclase (ppm). Data is grouped into lithofacies for ease of comparison.

Anorthosite Rb Sr Ba La Ce Pr Nd Sm Eu Gd Dy

07-1 Core 1.0 923 72 2.5 4.4 0.5 2.0 0.3 1.4 0.2 0.1

Rim 4.5 803 81 5.1 8.2 1.0 4.1 0.8 1.5 0.9 0.7

07-2 Core 0.5 904 98 5.1 7.1 0.7 3.1 0.3 1.7 0.2 0.1

Rim 1.8 830 94 6.1 8.8 0.9 3.8 0.3 1.6 0.2 0.1

21-1 Core 0.9 806 41 1.3 2.4 0.3 1.1 0.1 0.5 0.1 0.1

Rim 0.9 858 85 5.4 8.4 0.7 2.3 0.2 0.8 0.1 0.1

22-1 Core 0.5 760 84 6.3 10.4 1.0 3.9 0.5 0.9 0.3 0.2

Rim 6.2 903 71 5.4 7.5 0.7 2.5 0.2 0.8 0.1 0.0

22-5 Core 0.3 812 43 1.2 2.1 0.2 0.8 0.2 0.7 0.1 0.0

Rim 0.5 799 76 5.0 8.3 0.8 3.3 0.4 0.8 0.2 0.1

27-1 Core 0.8 788 52 4.1 7.9 0.8 3.2 0.5 0.7 0.2 0.1

Rim 0.8 805 68 4.3 6.8 0.6 2.4 0.3 0.9 0.2 0.1

Olivine Gabbro

32-1 Core 1.3 816 55 1.3 2.2 0.2 1.1 0.1 0.8 0.2 0.1

Rim 0.5 826 81 6.0 8.6 0.8 2.6 0.3 0.8 0.2 0.0

32-2 Core 8.0 805 85 1.8 2.9 0.2 0.9 0.1 0.7 0.1 0.1

Rim 0.8 836 72 2.7 4.7 0.5 2.1 0.3 0.9 0.2 0.1

32-3 Core 0.7 809 57 1.5 2.6 0.3 1.3 0.2 0.8 0.1 0.1

Rim 0.3 827 90 6.4 9.6 0.9 2.8 0.2 0.9 0.2 0.1

Troctolite

11-1 Core 1.4 721 40 1.3 2.7 0.4 1.6 0.2 0.7 0.2 0.1

Rim 1.1 816 60 3.5 7.0 0.8 3.0 0.5 0.9 0.3 0.1

13-1 Core 2.1 744 31 1.1 1.9 0.2 1.0 0.2 0.6 0.1 0.1

Rim 0.6 740 40 2.0 4.0 0.5 2.0 0.3 0.7 0.2 0.1

15-3 Core 0.5 617 35 2.1 3.9 0.4 2.3 0.2 0.5 0.2 0.1

Rim 0.9 801 43 3.2 5.4 0.6 2.5 0.2 0.8 0.3 0.1

15-6 Core 0.3 603 28 2.0 3.7 0.5 2.1 0.3 0.5 0.3 0.0

Rim 0.4 804 47 3.2 5.9 0.7 3.0 0.4 0.8 0.3 0.1

16b-2 Core 0.4 804 32 1.6 3.3 0.4 1.8 0.3 0.6 0.2 0.1

Rim 0.8 802 38 2.5 4.5 0.5 1.8 0.3 0.6 0.2 0.1

16b-4 Core 24.3 764 37 2.2 4.4 0.6 2.1 0.4 0.7 0.3 0.1

Rim 5.3 715 34 2.3 3.9 0.4 1.6 0.2 0.5 0.1 0.1

19-1 Core 0.8 812 43 1.1 2.3 0.3 1.0 0.1 0.6 0.1 0.1

Rim 1.0 877 41 4.8 8.7 0.8 2.4 0.4 0.8 0.2 0.1

Oxide Gabbro

35-3 Core 1.2 892 62 1.5 3.1 0.4 1.4 0.2 0.7 0.1 0.1

Rim 1.5 914 64 2.4 3.8 0.4 1.4 0.2 0.7 0.1 0.0

36-1 Core 1.2 911 62 1.4 2.1 0.2 0.9 0.2 0.8 0.1 0.0

Rim 3.0 971 70 2.8 4.0 0.4 1.8 0.2 0.8 0.1 0.1

36-2 Core 1.6 908 65 2.3 3.8 0.4 1.5 0.2 0.8 0.1 0.0

Rim 0.9 902 67 2.5 4.3 0.5 1.9 0.2 0.8 0.1 0.1

36-3 Core 1.9 917 67 2.0 3.5 0.4 1.3 0.1 0.7 0.1 0.0

Rim 93.1 855 95 3.3 5.0 0.4 1.8 0.2 0.8 0.2 0.1

36-4 Core 2.4 893 65 1.5 2.6 0.3 1.1 0.2 0.8 0.2 0.1

Rim 1.8 1011 74 4.4 6.8 0.6 2.3 0.3 0.8 0.2 0.0

38

Rare earth element spider diagrams

Rare earth element spidergrams for plagioclase grains, from core to rim, all share several defining

characteristics. When normalised to chondrites (McDonough and Sun 1995), plagioclase show light

rare earth element (LREE) enrichment, a pronounced positive Eu anomaly and heavy rare earth

element (HREE) depletion. Variation between individual grains is observed by changes in the slope of

the lines and by variation of absolute concentrations from core to rim. Plagioclase that exhibits

reverse zonation (Types 2 and 4) are characterised by a higher degree of variance in concentrations of

REE from core to rim (Figure 4.14a and 4.14b). While this characteristic is not unique to reversely

zoned plagioclase, all reversely zoned plagioclase have this high degree of variance. Normally zoned

plagioclase and plagioclase that show minor oscillatory zonation but are overall normally zoned

(Types 1 and 3) have limited variation and occasionally near-uniform REE compositions (Figure

4.14c and 4.14d).

39

Figure 4.14a (top) to 4.13d (bottom). REE spidergrams normalised to chondritic values (McDonough and Sun 1995)

depicting the typical variation between plagioclase types. Additional REE diagrams can be found in the electronic appendix.

40

Discussion

Mechanisms responsible for layering in a crystallising magma have evolved significantly since Wager

and Deer (1939) proposed fractional crystallisation and gravity settling to explain the layering

observed in the Skaergaard Intrusion. Magma recharge is often invoked to be the driver of layering in

mafic intrusions (Lombaard 1934; Irvine and Smith 1967; Tanner et al. 2014) since the addition of

new and more primitive magma to a chemically evolved liquid is the simplest explanation for altering

the effects of fractional crystallisation. Magma recharge also adds thermal energy into the system,

which can be localised at the base, driving convection cells in the magma chamber and facilitating the

transport of crystals in a viscous, evolving magma (Turner 1980). However, a better understanding of

magma chamber fluid dynamics demonstrates that the stratification of layers in a cooling liquid can

result from double diffusive convection without the addition of new magma (Irvine and Smith 1967;

McBirney and Noyes 1979; Irvine 1980; Huppert and Turner 1981; Huppert and Sparks 1984).

McBirney and Noyes (1979) contend that as a mafic magma evolves, it rapidly becomes non-

Newtonian in character and the ability of crystals to migrate through the chamber under the influence

of gravity is restricted. In a controversial paper, Marsh (2013) argues that physical processes in the

chamber such as gravity settling and flow alignment cannot be ignored as the effects are observable in

the textures of intrusive igneous rocks. Thus the exact role of magma recharge in the formation of

layered intrusion is of some debate in the literature and continues to be a contentious issue in modern

research.

McLeod, (1956, 1959b) proposed a model similar to Wager and Deer’s (1939) Skaergaard model to

explain the layering of the SDIC based on the presence of horizontally aligned feldspars and planar

interfaces between rock types. Unlike the Skaergaard Intrusion, the gross chemical variation of the

SDIC with height is inconsistent with a simple liquid line of descent. To account for this variation,

Mathison (1967, 1970, 1975, 1987), Riley (1991), and Walker (1998) combine magma recharge with

fractional crystallisation and gravity settling. Mathison (1967) proposed the injection and mixing of a

new magma into a cooling magma chamber as the major mechanism to explain the layering of the

SDIC (Figure 5.1).

41

Figure 5.1. A schematic model describing the evolution of the layers in the Somerset Dam taken from Mathison, 1967. (a)

New liquid is injected on top of partially crystallised semi-rigid layer via a feeder pipe in the floor. The new liquid convects

and mixes with the existing residual liquid. (b) Minerals are formed and sorted via gravity settling into layers (troctolite,

olivine gabbro, oxide gabbro and leucogabbro). (c) New magma enters the chamber, and some old liquid is ejected to

accommodate the new volume of liquid. Some crystals are entrained and ejected. The new layer created in (B) becomes fully

solid, and the cycle repeats.

To maintain a relatively constant magma chamber size, Mathison suggests that evolved liquids are

ejected when primitive magmas enter the chamber. Mathison’s interpretation is used by Riley (1991)

and Walker (1998) who view each cyclic sequence as a product of differentiation of a new pulse of

melt.

Understanding the role of magma recharge must take into consideration the fluid dynamics of

emplacement and magma interaction as well as the petrologic (and geochemical) evolution of the

magma. Over the past decade, a body of experimental and theoretical work has been developed to

evaluate the physical aspects of magma intrusion (Cruden and McCaffrey 2001; Bunger and Cruden

2011; Annen 2011; de Saint Blanquat et al. 2011; Menand 2011; Zibra et al. 2014). This work

clarifies the importance of the size and frequency of magmatic pulses relative to chemical

composition – its application to field studies is ongoing. The petrologic and geochemical evolution of

magma is recorded in the minerals and rocks that comprise the intrusion. This record may be affected

by numerous processes, ranging from subsolidus alteration and metasomatism (Irvine 1980; Holness

et al. 2006) to thermal re-equilibration (Tanner et al. 2014).

42

Thermal Re-equilibration

Subsolidus re-equilibration of plutonic minerals can significantly alter the original magmatic chemical

profiles of individual minerals. There is extensive evidence for diffusive modification in large and

long-lived intrusive bodies (e.g., The Bushveld Complex; Tanner et al., 2014). Heat generated by the

crystallisation in combination with the heat stored in a large volume magma chamber sustains a heat

anomaly post crystallisation and facilitate thermal diffusion. A chemically zoned crystal that has a

substantial chemical potential difference from core to rim will attempt to reach equilibrium over time

and become homogeneous. Subsolidus diffusion is a function of time, temperature, physical properties

of the cations, and the magnitude of the potential difference in concentration gradients, oxygen

fugacity and partition coefficients. Small radius and low charge cations diffuse more rapidly than

those cations that are larger in size and higher in charge (Brady and Cherniak 2010).

Olivine, pyroxene and plagioclase are the major minerals that comprise the layers of a layered mafic

intrusion like the SDIC. The effect of subsolidus equilibration on these minerals can compromise or in

the worst cases obscure the subtle chemical signatures of magma recharge. Due to olivine’s structure

of isolated SiO4 tetrahedra linked by an octahedral site, the M1 site only forms a chain in one plane

(001) while the M2 site is not joined in any direction (Chakraborty 2010). This results in the simple

binary substitution between the dominant cations (Mg and Fe) in olivine diffusing more rapidly and

with a lower activation energy (Brady and Cherniak 2010; Chakraborty 2010; Zhang 2010). In

contrast, the more highly polymerised tectosilicate structure of plagioclase combined with the more

complex coupled substitution of Ca-Al and Na-Si in plagioclase results in relatively slow diffusion.

Pyroxene, a modestly polymerised chain silicate, but possessing more complex chemical

substitutions, has intermediate diffusion rates (Müller et al. 2013).

To constrain the effect of subsolidus diffusion in the minerals of the SDIC, some calculations were

achieved utilising published data (Grove et al. 1984; Chakraborty 1997; Müller et al. 2013).

Diffusivities for the mineral species were obtained using the Arrhenius equation:

= /( )

43

Where A is the pre-exponential factor and Ea is the energy of activation. These diffusivities were then

used with the diffusion equation to obtain finite distances of diffusion:

=√4

Where k is the diffusivity obtained from use of the Arrhenius equation and t is time. Because a

magma chamber cools over time it is a non steady state diffusion problem. To provide an estimate of

the ability of minerals of the SDIC to undergo diffusive modification this model was solved for

temperatures between an assumed solidus of 900o and an assumed final temperature of 400o in 10o

increments. This was done for a timescale of 30Ka, 10Ka and 1Ka (Figure 5.2) and show that the

diffusivity length scales between olivine and plagioclase differs by as much as 4 orders of magnitude.

These timescales were chosen as studies suggest that cooling can be achieved in as little as ~51 years

in mafic sills to ~52,000 years in moderately sized epizonal felsic plutons (Nabelek et al. 2012).

The majority of diffusion within the crystal is achieved early on in the subsolidus cooling period.

Plagioclase has a maximum potential diffusion length of ~10µm ceases to diffuse in meaningful

amounts (<0.1µm) below approximately 750Co while olivine has a maximum potential in excess of

1000µm (1mm) and continues to diffuse until the lower bound of 400Co is reached. Thus olivine has a

much greater potential for chemical diffusion and is reasonable to assume that olivine crystals from

the SDIC have experienced significant subsolidus diffusive modification. Given the average crystal

sizes for olivine (<4mm) and plagioclase (1-3mm) in the SDIC as well as its short cooling history,

SDIC olivine is thought to be strongly re-equilibrated, clinopyroxene is modestly re-equilibrated, and

plagioclase is not re-equilibrated. This evidence combined with the calculated diffusion rates (Figure

5.2) implies that plagioclase retains its crystallisation history and can be used as an accurate record of

chemical change in the magma body.

44

Figure 5.2. Calculated diffusivity length scales assuming a temperature of 900oC to 400oC and a crystallisation periods of

30ka, 10ka and 1ka. Lines plotted on these graphs represent the solution to the diffusion equation at a steady state

temperature given on the x axis for a time period indicated by the legend. The left figure highlights the rapid early diffusion

while temperatures are high. The right figure is logarithmic to give a clearer impression of the magnitude of the diffusion

length scales.

Controls on the chemical profiles of plagioclase and olivine

Variations in plagioclase major element compositions can often be attributed to changes in

temperature, pressure, water content and magma composition (Smith, 1972). Complex major element

zonation of plagioclase can, therefore, have multiple origins. However, it is unlikely that the SDIC

experienced dramatic changes in pressure during crystallisation. It is reasonable to consider the

magma to be water poor based on the absence of common primary hydrous minerals and a low

abundance of fluid inclusions in minerals (Walker 1998). The most reasonable explanation for

plagioclase compositional variation in the SDIC must be a combination of changes in temperature and

magma composition.

A similar argument can be made for the effect of environmental conditions on the minor and trace

element composition of plagioclase. Highly incompatible trace elements in particular (such as REEs)

are unaffected as these concentrations are controlled by the composition of the melt; their partition

coefficients are relatively insensitive to these environmental changes (Ginibre et al. 2002, 2007;

45

Ginibre and Wörner 2007; Niu et al. 2014). Thus, the minor and trace element compositions of

plagioclase are indicative of melt chemistry.

The variation in rock and mineral chemistry in the examined portion of the SDIC is attributed to

magmatic differentiation as a result of fractional crystallisation (McLeod 1959b). Plagioclase in the

stratigraphy is typically normally zoned as shown by plagioclase core to rim variation with height

(Figure 4.3) and is an adcumulate phase. In-situ growth reflects the interaction of the inter-cumulus

liquid with the main body magma (Irvine 1980) as the evolving liquids in the crystallisation front are

ejected from pore spaces during compaction and ascend and mix with the main body of liquid (Tait et

al. 1984; Morse 1986; Tait 1996). Some of the normally zoned plagioclase observed can be attributed

to cooling and chemical differentiation of the magma body abetted by this process. Other normally

zoned plagioclase have experienced a rapid change in magma chemistry which is reflected in a sharp

change in mineral chemistry (e.g., Figure 4.1c, 4.8d and 4.12d).

Variations in phase chemistry and rock texture

Major element and textural variations

The chemical evolution of plagioclase as a function of position in the SDIC (Figure 4.3) poses some

problems for a simple fractional crystallisation model. At the base of the sampled stratigraphy, the

composition of the plagioclase is relatively evolved (An55). Moving up stratigraphy, the median and

spread of compositions increase in anorthite component to a maximum of An85 at 300m in height.

The composition subsequently declines towards the top of the sampling sequence (An55). The most

primitive unit in the stratigraphy is the large central troctolite unit between ~245m and ~325m and is

comprised of samples SD11, 13, 15, 16, and 19.

The plagioclase in the central troctolite are a contrast to the rest of the stratigraphy due to their high

anorthite content but also are diverse due to variations in grain size and orientation. The stratigraphic

column (Figure 4.3) in conjunction with photomicrographs of the thin section textures (Figure 4.6)

highlight this variation. At the base of and top of the central troctolite, the rock textures are coarser

grained with no observable grain alignment and are typically adcumulates. Towards the centre of the

46

unit the grain size becomes finer, is strongly aligned along the c axis, and has both mesocumulate and

adcumulate textures within the samples. The change in texture requires a localised mechanism for

aligning the grains. If interpreted as evidence of fluid transport, this suggests the addition and

transport of new magma at this horizon in the stratigraphy. The new magma propagates laterally

through the sill through a crystalline mush where the previous magma was fractionating. As the sill

increases in thickness due to the addition of new liquid intraplating the existing sill, the crystallisation

front is pushed apart. Crystals in cumulates that form SD 11 and SD 19 (the base and the top of the

central troctolite) might have at one time been formed in the same mush but subsequently pushed

apart by the addition of new liquid and the expansion of the sill to accommodate the additional

volume.

True oscillatory zoning and reverse zoning is rare in plagioclase in the SDIC when considering the

stratigraphy as a whole. Where these features are present, they are concentrated in an approximately

30 meters thick zone located in the central troctolite unit between 270m and 300m. Plagioclase

(An55-An60) in samples from this zone (SD 13 and 15) are commonly oscillatory zoned, with an

overall reversed zoned character (Figure 4.7, Table 4.6). They are characterised by reabsorption

textures, particularly in SD15. In contrast, sample SD 16 which lies immediately above SD 15 (within

~10m) exhibits mesocumulate textures with plagioclase that is normally zoned, with markedly higher

anorthite content (up to An80-85). This contrast in anorthite composition in combination with textural

features suggests that the stratigraphic horizon between 270-300m experienced a disturbance from a

typical liquid line of descent during its crystallisation history. Other authors have used similar

evidence as a basis for identifying magma recharge in plutonic rocks (Streck 2008). This disturbed

liquid line of descent in the central troctolite could be a zone of magma recharge, where a hotter, more

primitive magma interacted with a cooler and more differentiated magma.

Olivine trends are similar to that of plagioclase, although significantly suppressed. Olivine is more Fe-

rich at the base and top of the sampled stratigraphy and more Mg-rich towards the middle of the large

centralised troctolite unit (Table 4.5). Olivine is, however, much less variable in composition both

47

within a single crystal and as a function of height (typically less than 1% Fo). This homogeneity is

attributed to subsolidus re-equilibration.

Minor and trace element variations

As discussed above (see ‘Controls on the chemical profiles of plagioclase and olivine’), the minor and

trace element compositions of minerals are a product of a particular element’s concentration in the

melt and the liquid/mineral partition coefficient. The chemical profile of minor and trace elements in a

mineral can offer evidence of a change in magma chemistry, (Ginibre and Wörner 2007). Bulk

composition change of the magma is the most dominant control on the variation of trace elements in

plagioclase. In this study, minor and trace element data from plagioclase determined by LA-ICP-MS

is collected at a relatively coarse spatial resolution (45µm) for each ablated ‘spot’. This data

represents an average of a more complex fine scale zonation patterns detected by EPMA (10µm spot).

Zones of plagioclase that record a change in major element chemistry (e.g. an increase or decrease in

Ca content resulting in a change in An component) will also record changes in minor and trace

element chemistry dependent on their partition coefficients (Blundy and Wood 2003). For example,

the relative compatibility of Ba in plagioclase is dependent on the Na concentration in the mineral

since Ba is more similar to Na than Ca. The opposite is true for Sr and Fe which are more similar to

Ca (when Fe is in its 2+ valence state) (Ginibre et al. 2007). In the SDIC Fe and Sr concentrations

increase with An component while Ba concentrations decrease. The relative increase of Sr and Fe to

An component is significantly muted. Thus a significant increase in An component might only result

in a small increase in Sr or Fe concentrations due to the overall incompatible nature of these elements.

While Fe has an ability to diffuse out of plagioclase (Tanner et al. 2014) in plutonic systems, the

SDIC is small enough and cooled rapidly enough to limit the effects of subsolidus diffusion (see

above section ‘Thermal Re-equilibration”).

Plagioclase of the central troctolite unit (samples SD15, 16 and 19; Figure 4.13) are characterised by

variations in Ba, Fe and Sr concentration that are different from the rest of the sampled SDIC

plagioclase. Despite significant An variation in SD 15 (Figure 4.12a and b), there is little change in

48

the minor element concentrations. Since the bulk composition of the host magma has a much stronger

impact on the concentrations of Sr, Fe and Ba in the plagioclase than a change in partition coefficient

as a result of changing An component in the crystal (Ginibre et al. 2002; Ginibre and Wörner 2007) a

possible explanation for this anomalous trace element concentrations could be chemical buffering by

the addition of new liquids.

Partition coefficients for rare earth elements (REE) in plagioclase indicate that these elements are

highly incompatible with plagioclase. As such, the concentration of rare earth elements in a normally

zoned plagioclase will be greater in the rim relative to the core if pure fractional crystallisation has

occurred. REE chondrite-normalised spider diagrams from a plagioclase that is normally zoned

(Figure 4.14c and d) is consistent with this process. This is consistent with previous REE data analysis

done by Walker (1998) which demonstrated that REE patterns plagioclase in the SDIC are typically

consistent with what is predicted by fractional crystallisation. However, this is not the case for

plagioclase from the central troctolite unit that are reversely zoned have more complex histories with

more variation in REE concentrations from core to rim (Figure 4.13a and b).

Composite diagrams with characteristic REE patterns for the samples from the central troctolite where

recharge is inferred are shown along with their corresponding An variation (core to rim) and BSE

images (Figure 5.3). These REE diagrams, normalised to the core from each sample, have arrows that

indicate the relative enrichment and depletion of REE from core to rim. Plagioclase from SD16 have

anorthite-rich cores (An80-85) and are normally zoned - often with a steep drop in An content near

the rim (An65-70) (e.g., Figure 4.12c). These plagioclase also show a consistent enrichment in REE

from core to rim (Figure 5.3a). REE contents are correlated with An content in that a decrease in An

content is mirrored by an increase in REE content. In rocks where reverse zoning in plagioclase is

common (such as SD15, Figure 5.3b), the REE patterns suggest a more complex crystallisation

history. In SD15 plagioclase 3 (Figure 5.3b), REE concentrations become depleted relative to the core

concentrations for a zone approximately 96µm. The decrease in REE concentrations is matched by a

corresponding enrichment in An component. The next zone (from 96µ to the rim) steadily decreases

in An component and enriches in REE elements. This pattern indicates a crystallisation history where

49

the plagioclase core was exposed to a more primitive liquid, disturbing the predicted liquid line of

descent.

Figure 5.3a (left stack) and b (right stack). Rare earth element spidergrams along with An component variation relative to the

core and corresponding BSE image. The rare earth element spidergrams are normalised to the core composition of each

sample. Arrows indicate enrichment and depletion of REE relative to the core. Since these samples are core normalised, the

core is plotted along the 100 line for each REE diagram.

50

What constitutes a ‘macrocycle’?

Previous interpretations of the SDIC propose that the stratigraphy is comprised of seven macrocycles

that progress from leucogabbro to troctolite, to an olivine gabbro and finally an oxide gabbro, and that

each macro-cycle is the product of magma replenishment (Mathison 1987). The observed plagioclase

compositions and zonation patterns in this study present some inconsistencies with this model.

First, if each new cycle represents the injection of a new and more primitive magma into the magma

chamber, then plagioclase at the interfaces between the macrocycles (postulated boundaries between

existing and new magma batches) should contain a marked increase in anorthite component (i.e.

reverse zoning). The samples in this study collected across two of the proposed ‘contacts’ between the

macrocycles (Figure 2.1 and Figure 2.2) indicate that the composition of the plagioclase grains is

variable and that the plagioclase with the highest median An component does not occur at these

contacts (Figure 4.3). Instead, plagioclase located at these boundaries have lower An content than the

plagioclase sampled from the centre of a macrocycle unit.

Second, plagioclase with reverse zoning commonly occurs within the central troctolite of this study

area (Table 4.2). These plagioclase are also characterised by reabsorption textures of plagioclase cores

(Figure 4.7). These characteristics are interpreted to be the result of magma mixing during

crystallisation. This evidence can be used to understand the evolution of the SDIC in different ways.

In the existing interpretation of the SDIC, each macrocycle is the result of a discrete pulse of liquid

entering a single magma chamber and interrupting the liquid line of descent (Figure 5.1). In this

scenario, the anomalous third macrocycle (3A and 3B, Figure 2.2) is emplaced and crystallises but is

interrupted by the addition of new liquid before becoming solid, but after some differentiation has

been done.

If one considers that the chemical evolution of the third macrocycle is disturbed by magma recharge

then macrocycles 3 and 4 could be speculated to be the differentiation of a single sill. Following this

line of reasoning, the original macrocycle encompassed all material between 250m and 400m and was

undergoing fractionation until the recharge event occurred at ~300m disturbing the chemical

51

evolution of the stratigraphy and being recorded in the rock record between ~250m and ~350m.

According to Mathison’s (1987) interpretation (Figure 2.2), each macrocycle is reversely fractionated

from leucogabbro to troctolite then normally fractionated from troctolite to oxide gabbro. This is

based on the work of other authors that suggest that cryptic reversals are delayed and gradual rather

than abrupt (e.g.,. Dunham and Wadsworth, 1978), and the assumption that the leucogabbro is the

base of the sequence. Absent the proposed recharge event in the central troctolite, treating

macrocycles 3 and 4 as a single macrocycle fits with the observed evolutionary trends of the other

macrocycles (or sills). (Figure 5.4).

This hypothesis is supported by the distribution of trace elements in the sampled stratigraphy. The

concentration of Eu in the rocks of the central troctolite and above share many similarities (Figures

4.9, 4.10 and 4.11). These samples comprise the interrupted macrocycle 3 and macrocycle 4 in

Mathison’s interpretation. The partition coefficient of Eu in plagioclase varies only slightly with

changing An component (Figure 8 in Bindeman et al., 1998). Thus, the concentration of Eu in the

rocks interpreted as macrocycle 3 and 4 plot consistently between 0.5 and 1.0 ppm regardless of

changes in An component (Figure 4.9). Samples that were taken from below the large central

troctolite unit (interpreted by Mathison as macrocycle 2) are characterised by distinctly different

concentrations of Eu, typically between 1.0 and 2.0 ppm. This is not explainable by simple fractional

crystallisation and implies that the rocks below the central troctolite unit crystallised were in

equilibrium with a different magma than the rocks sampled from within the central troctolite and

above.

The evidence discussed in this paper is not sufficient to offer a definitive answer to this hypothesis.

Regardless, there is strong evidence in the phase chemistry, and mineral textures that indicate a

potential recharge event occurred within the large central troctolite unit, rather than at either of the

two previously interpreted contacts between macrocycles.

52

Figure 5.4. Variation in An component with stratigraphic height modified after Mathison (1987). The blue circle highlights a

significant change in An component which is proposed to be the result of magma recharge. The red line indicates a

speculative chemical evolution of the sill absent magma recharge. Lithologies are shown (white = leucogabbro, dots =

troctolite, dashes = olivine gabbro, black = oxide gabbro). Macrocycles after Mathison (1987) are shown on the right. The

sampling area for this study is indicated by a dotted line to the right of the stratigraphy.

53

Composition of ‘recharge’ magma

If magma recharge played a role in the evolution of the SDIC, then it cannot be assumed that all the

pulses of liquid share the same parental source. To determine if any recharge events tapped the same

source material as the original liquid a model of liquid compositions for plagioclase has been made.

Plagioclase trace element chemistry is used to calculate the composition of the equilibrium liquid for

each stage of plagioclase growth. Partition coefficients were calculated using the method of Bindeman

et al., 1998 (Table 5.1).

Table 5.1 – Calculated partition coefficients after Bindeman et al. (1998) for plagioclase/basalt at 1523o K.

Fe Rb Sr Ba La Ce Pr Nd Sm Eu

< An55 Average 0.30 0.04 3.45 0.47 0.19 0.06 0.15 0.17 0.15 0.76 Minimum 0.25 0.03 2.98 0.36 0.18 0.05 0.13 0.15 0.13 0.76 Maximum 0.47 0.07 5.07 0.94 0.22 0.10 0.20 0.22 0.21 0.76

An55 - An65 Average 0.21 0.03 2.56 0.27 0.17 0.04 0.12 0.14 0.12 0.76 Minimum 0.18 0.02 2.25 0.21 0.17 0.03 0.11 0.13 0.10 0.76 Maximum 0.25 0.03 2.97 0.36 0.18 0.04 0.13 0.15 0.13 0.76

An65 - 75Average 0.16 0.02 2.01 0.18 0.16 0.03 0.10 0.12 0.10 0.76 Minimum 0.13 0.02 1.71 0.13 0.15 0.02 0.09 0.11 0.08 0.76 Maximum 0.18 0.02 2.24 0.21 0.17 0.03 0.11 0.13 0.10 0.76

> An75 Average 0.12 0.01 1.59 0.12 0.15 0.02 0.08 0.10 0.08 0.76 Minimum 0.11 0.01 1.45 0.10 0.14 0.02 0.08 0.10 0.07 0.76 Maximum 0.13 0.02 1.69 0.13 0.15 0.02 0.09 0.11 0.08 0.76

The calculated liquid compositions are broadly similar with all liquids being enriched in LREE

relative to chondrites (Figure 5.5). A selection of minimum and maximum core and rim samples is

presented in Table 5.2. The absolute size of the positive Eu anomaly depends on the partition

coefficient, which is affected by the oxidation state of the liquid. In most terrestrial magmas Eu exists

more commonly in a trivalent state along with some amount of divalent Eu depending on the redox

conditions of the melt (Schreiber 1977). This study uses a Kd of 0.76, which falls midway into the

range of Kd for Eu (0.1 – 1.5) tabulated by Rollinson (1993). The positive Eu anomaly of the

calculated liquid implies that this liquid was enriched in Eu. Other authors have noted anomalous

54

enrichment of Eu in basaltic liquids (e.g., Mazzucchelli et al., 1992) with a possible cause being

mixing between mantle-derived and crustal-derived melts.

In normally zoned plagioclase the modelled liquid enriches progressively in REE elements. The

liquids which are in equilibrium with the rims of the crystals may be more enriched in REE elements

by as much as a factor of 5 when compared to the core (Figure 5.5 e, f and g). This trace element

pattern is consistent with what is predicted by fractional crystallisation. However, despite being

normally zoned, plagioclase from SD 16 (within the central troctolite, Figure 5.5d) has a REE

evolution from core to rim that is quite similar to SD 15 (Figure 5.5c) which is a reversely zoned

plagioclase. Both SD 15 and SD 16 become depleted in REE when moving away from the core and

become enriched again back to core-like values at the rim. This is inconsistent with what is expected

from a normally zoned plagioclase and might indicate the interaction between different liquids as a

mechanism for disturbing the liquid compositions from their predicted liquid line of descent. The

overall character of the REE spidergrams suggests that if a recharge event took place in the SDIC that

the new and pre-existing liquids would have originated from the same source.

55

Figure 5.5. Trace elem

ent composition of the m

odelled original liquid from the S

omerset D

am layered series. T

he samples have been arranged in stratigraphic order from

A

at the bottom to H

at the top. Norm

alised to chondrites after Sun and M

cDonough (1989)

56

Table 5.2. Modelled trace elements for core and rim analyses from each sample. The minimum and maximum core and rim values have been selected and are not necessarily from the same crystals.

Sample An% Rb Sr Ba La Ce Pr Nd Sm Eu Gd Dy

2 46 core 3 996 83 5.2 8.1 0.8 2.8 0.3 1.2 0.2 0.1

60 core 1 922 50 2.4 4.2 0.5 2.0 0.4 1.4 0.3 0.2

43 rim 1 896 198 9.5 12.0 1.2 3.7 0.4 2.1 0.4 0.0

62 rim 0 923 70 2.3 4.2 0.6 2.0 0.2 1.5 0.3 0.1

7 51 core 0 904 98 5.1 7.1 0.7 3.1 0.3 1.7 0.2 0.1

58 core 1 869 45 1.9 2.8 0.3 1.2 0.1 0.7 0.1 0.0

40 rim 4 803 81 5.1 8.2 1.0 4.1 0.8 1.5 0.9 0.7

57 rim 0 868 46 1.9 3.2 0.3 1.6 0.2 0.9 0.1 0.0

11 63 core 1 721 40 1.3 2.7 0.4 1.6 0.2 0.7 0.2 0.1

77 core 2 779 36 1.2 2.6 0.3 1.3 0.2 0.8 0.1 0.1

61 rim 2 726 43 1.6 3.3 0.4 1.6 0.3 0.8 0.2 0.1

76 rim 1 841 54 3.3 7.1 0.8 3.1 0.5 0.9 0.3 0.2

13 44 core 6 680 135 5.6 6.7 0.5 1.4 0.1 0.7 0.1 0.0

63 core 2 744 31 1.1 1.9 0.2 1.0 0.2 0.6 0.1 0.1

59 rim 0 727 51 3.0 5.3 0.5 2.1 0.3 0.6 0.1 0.0

67 rim 1 605 48 2.5 5.6 0.7 3.8 0.9 0.8 0.8 0.9

15 55 core 1 683 40 1.3 2.4 0.3 0.9 0.1 0.5 0.0 0.0

69 core 1 688 31 1.3 2.6 0.3 1.3 0.2 0.6 0.1 0.1

70 rim 1 801 43 3.2 5.4 0.6 2.5 0.2 0.8 0.3 0.1

75 rim 0 793 34 2.6 4.2 0.5 1.6 0.2 0.7 0.2 0.0

16 72 core 1 837 39 1.8 3.6 0.5 1.9 0.3 0.7 0.2 0.1

80 core 24 764 37 2.2 4.4 0.6 2.1 0.4 0.7 0.3 0.1

57 rim 5 715 34 2.3 3.9 0.4 1.6 0.2 0.5 0.1 0.1

73 rim 2 799 31 1.3 2.7 0.3 1.2 0.2 0.6 0.2 0.1

19 62 core 1 898 46 1.1 2.1 0.3 1.0 0.2 0.7 0.1 0.1

68 core 1 860 38 1.2 2.2 0.3 1.1 0.2 0.6 0.1 0.1

60 rim 12 837 81 5.0 7.8 0.7 2.6 0.4 0.8 0.1 0.1

64 rim 2 547 36 2.7 6.5 1.0 4.9 1.4 0.7 1.6 1.7

21 61 core 3 824 53 1.9 3.6 0.4 1.6 0.3 0.8 0.2 0.1

65 core 1 819 39 1.1 2.0 0.2 0.8 0.1 0.6 0.1 0.0

53 rim 1 689 40 1.2 2.4 0.3 1.2 0.2 0.6 0.2 0.1

58 rim 1 799 72 5.1 8.2 0.8 2.7 0.4 0.7 0.2 0.1

27 59 core 3 753 42 4.6 8.2 0.9 4.2 0.6 0.9 0.4 0.3

73 core 1 750 33 1.5 2.9 0.3 1.4 0.2 0.6 0.2 0.1

60 rim 1 805 68 4.3 6.8 0.6 2.4 0.3 0.9 0.2 0.1

66 rim 4 830 69 3.7 6.3 0.7 2.6 0.4 0.8 0.2 0.1

32 48 core 8 805 85 1.8 2.9 0.2 0.9 0.1 0.7 0.1 0.1

62 core 1 809 57 1.5 2.6 0.3 1.3 0.2 0.8 0.1 0.1

51 rim 0 827 90 6.4 9.6 0.9 2.8 0.2 0.9 0.2 0.1

57 rim 1 836 72 2.7 4.7 0.5 2.1 0.3 0.9 0.2 0.1

36 56 core 4 876 73 2.2 4.2 0.4 1.9 0.3 0.8 0.2 0.1

57 core 2 917 67 2.0 3.5 0.4 1.3 0.1 0.7 0.1 0.0

52 rim 93 855 95 3.3 5.0 0.4 1.8 0.2 0.8 0.2 0.1

58 rim 3 971 70 2.8 4.0 0.4 1.8 0.2 0.8 0.1 0.1

57

SDIC – Revised Model

An early model for the formation of the SDIC (Figure 5.1) involves a single magma chamber which

undergoes differentiation and forms a rigid or semi-rigid layer at its base (Mathison, 1967). A similar

model was proposed by G.M. Brown to explain the occurance of laying observed for the Rhum

Intrusion in the Hebrides (Brown 1956). In this model, new magma periodically enters the chamber

from a feeder dyke at its base, penetrating the existing solid to semi-solid layers already crystallised

on the floor of the chamber, and mixing with the residual liquid. Some residual liquid is ejected from

the top of the chamber to create space for the replenishment of new primitive magma from below.

This process results in a sequence of layers (macrocycles) stacked upon each other comprising a

sequence of rocks (leucogabbros, troctolite, olivine gabbro, oxide gabbro). In this model, the strongest

evidence of magma recharge and liquid interaction should be located at the interface between the

macrocycles. In this study, samples across two such interfaces (between macrocycles 2 and 3 and

again between 3 and 4) reveal no easily detectable sign of magma recharge. Instead, the strongest

evidence is located in the middle of the large, central troctolite unit in macrocycle 3.

The SDIC is approximately circular in shape, 3km across and 500m in thickness (Walker 1998) with

gently inward sloping contacts resulting in a total volume of approximately 3.5km3 and a length to

thickness ratio of 6:1. Taken as a single magma body, this geometry and length to thickness ratio are

similar in size to a laccolith (Bunger and Cruden 2011). However, the SDIC is unlike a laccolith in

that is lacks the defining characteristics of steep sides and up-bending of wall rocks, and the nature of

the upper and lower contacts are uncertain. The aspect ratio of the SDIC argues against it being a sill

unless it is comprised of a number of sills, closely related in time and space. Field work that formed

the basis of previous studies notes sharp contacts between the macrocycles (Mathison 1967; Walker

1998). If each macrocycle is interpreted as an individual sill with a thickness of 10m to 100m then,

the corresponding aspect ratios range from 30:1 to 300:1- very much sill-like in dimensions (Bunger

and Cruden, 2011). Thus, a model for the SDIC is that it is a composite intrusion, formed from the

sequential emplacement of a number of sills being similar in composition.

58

Sill emplacement occurs primarily at existing interfaces between layers of contrasting mechanical

properties due to neutral buoyancy, rotational stresses, rigid anisotropy in the country rock and

contrasting rheology between the magma and the country rock (Menand 2011). Contrasts in rigidity

between materials provide the ideal emplacement opportunity for sills, and can result in subsequent

sills intruding along planes of rigid contrast created by an earlier sill and the country rock (Kavanagh

et al. 2006). The emplacement of the sill results in fracture propagation along the minimum axis of

compressive stress, σ3, which is assumed to be horizontal (Burchardt 2008). After the initial sill is

formed additional magma pulses ascend and will either underplate, overplate or intraplate the existing

sill(s) depending on the sill’s degree of crystallisation, rigidity, and the density of the new liquid

(Annen 2011; Miller et al. 2011).

As a composite differentiated sill grows, the volume and frequency of each successive injection of

magma and their relative locations impact the solidification timescale of each successive batch of

liquid. The time needed for each magma batch to cool and equilibrate with the country rock is long

relative to the potential time between magma pulses (Annen 2011) which can be as little as 10s of

years. The volume of the SDIC is similar to that of the Black Mesa Sill and suggests an emplacement

timescale of ~5-10 years between pulses with total emplacement taking no more than 1,000 years with

an emplacement rate of ~0.01km3yr-1 (Figures 3 and 4 found in de Saint Blanquat et al., 2011).

As a system evolves, the thermal contrast between the new hot magma and the country rock grows

smaller since the country rock is already heated to a degree by previous intrusions. Huppert and

Sparks (1988) demonstrated that despite the relatively short lifespans basaltic sills (e.g. 1.5km cools in

270 years), these intrusions could still generate significant amounts of silicic melt. Repeated

intrusions increase the ambient temperature of the crust, increasing the volume of crustal melting and

slowing the cooling of the intrusions. Huppert and Sparks (1988) note that their work is applicable

only to the deep crust, where ambient temperatures are 500oC or greater. Extensive modelling (Annen

2005, 2011) demonstrates that when the system is thermally mature, composite sills can experience

limited remelting and obscure the contacts between the different intrusive events. Thus, a composite

sill may have an appearance of being formed by a single, much larger volume emplacement event.

59

It is clear from numerous field studies that sills of varying compositions and sizes are emplaced into

the shallow crust and that many of these show a complex interplay between the incremental nature of

magma emplacement and plutonic differentiation (Coleman et al. 2004; Glazner et al. 2004). To

facilitate emplacement in the crust, many plutonic bodies may be the product of the amalgamation of

multiple separate pulses of liquid - the contacts between these pulses can often be subtle and difficult

to identify in the field (e.g. Miller et al., 2011). ‘Composite’ intrusions may, in fact, be the normal

case (Menand 2008).

A new model for the SDIC builds upon the original ideas of Mathison (1967) who recognised the

likelihood of periodic magma replenishment to create the various macrocycles. The new model

proposes that the SDIC is a composite intrusion created from the sequential emplacement of mafic

sills over a short time in the same area. Unlike the previous model, the interaction between each sill is

limited as the new magma either under plates or over plates the existing mostly solid (and rigid) sill.

As more magma is emplaced into the region, ambient temperature rises, limited melting occurs along

sill contacts, and ‘diffuse’ boundaries are created between different sills (Figure 5.6). Magma

mixing/mingling does not occur in most cases. The exception to this general scenario is the third

macrocycle (Figure 2.2) which contains the large central troctolite unit.

The third macrocycle, considered to be an ‘interrupted’ macrocycle by Mathison (1987), is the

product of sill intraplating in which relatively crystal-poor new magma is emplaced into the

crystalline mush of a previously emplaced sill (see Figure 8, Miller et al., 2011). As a sill crystallises

from the base and roof, a new magma that is injected into the system rises until it contacts the semi-

rigid ‘trap’ at the roof of the growing sill. The new melt and any entrained crystals then propagate

laterally and mixes with the existing mush (Figure 5.6). As the crystals from the new and old liquid

are exposed to the change in both melt chemistry and temperature, they develop disequilibrium

features, such as reverse and oscillatory zoning and reabsorption textures in the cores of the

plagioclase. Crystals from the primitive, newer magma (primocrysts after Marsh, 2013) which might

have a high anorthite component come into contact with a more evolved liquid and become normally

zoned. If the difference in the liquid compositions is sufficient, then the difference in the An

60

composition of the various zones might be sharp rather than a gradual enrichment of Na as predicted

by a simple model of fractional crystallisation. All of these characteristics are observed in the

plagioclase sampled from the central troctolite unit which forms the majority of the so-called

‘interrupted’ third macrocycle (e.g., Figure 4.13 c and d).

It is unlikely that significant convection occurs to homogenise the new magma with the residual

liquid, however mixing may be possible as the new magma propagates laterally through the existing

sill. The expected internal variations under this model will vary greatly depending on where the layers

are sampled and on the degree of propagation and mixing achieved by the new liquid. Absent whole

scale convection to homogenise the new magma mixture, the heterogeneous nature of the crystalline

rock as it cools will provide a stronger or more muted evidence of its crystallisation history depending

on where the samples were collected. Simply put, sampling a drill hole that is located proximal to the

feeding dikes of the sill might have very different geology and geochemical profiles than sampling a

drill hole that is distal to the pipe entrance.

The order in which the sills are emplaced is unknown but can be speculated. The numbering of the

macrocycles from 1 through 7 is for convenience relative to stratigraphic height and in no way implies

an order of emplacement. Mathison (1987) notes that macrocycles 1 and 2 are significantly more

evolved than macrocycles 3 and 4. If all the sills forming the SDIC are sourced from the same

parental magma, as is implied from the liquid modelling in this study, then macrocycles 1 and 2 might

represent a late stage sill that is the product of more evolved liquid ascending from whichever source

is feeding the system and underplate macrocycle 3. Because the sills can be emplaced relatively

quickly obtaining exact dates for each sill and thus the quantitative order of emplacement is unknown.

61

Figure 5.6 Model of sill accumulation and magmatic recharge for the SDIC. 1. The Neera Volcanics prior to the

emplacement of the first sill. The dotted line depicts a planar feature such as a bedding plane that is ideal for the

emplacement of a new sill. 2. The first pulse of liquid ascends and is emplaced along the existing lateral boundary between

the bedded strata, cools and becomes solid or near solid prior to the emplacement of the next magma 3. A new pulse of

liquid ascends through the crust and contacts the new planar interface created by the first sill. The magma can either

underplate (as shown in 3) or overplate (as shown in 4) depending on density differences and mechanical rock properties.

Once emplaced this magma will also cool and crystallises 5. Should additional melt ascend prior to each sill fully

crystallising it may intraplate and mix with the pre-existing sill. The addition of new magma into the crystallising sill

62

disturbs the fractional crystallisation process and creates various disequilibrium features in the forming minerals. 6.

Subsequent sills are added forming the seven macro-cycles observed in the SDIC. Note that these ‘macrocycles’ are labelled

after Mathison (1 through 7) and do not in any way imply their relative order of emplacement. 7. A more detailed

examination of magmatic intraplating of a crystallising sill described in (5). As the new liquid propagates laterally through

the sill existing crystals that are entrained by the new liquid come into contact with the move evolved liquid in the sill, and

minerals crystallising in the sill come into contact with the more primitive liquid injected into the crystallisation front.

7a Plagioclase crystal located in the old crystallisation front is exposed to the new liquid disturbing its natural fractionation

and resulting in the crystal becoming reversely zoned. 7b Entrained plagioclase that is brought into the chamber with the

new magma. Upon contacting the more evolved older liquid, the crystal becomes normally zoned with a sharp contrast

between the older more An-rich core and the new An poor rim.

63

6. CONCLUSION

Detailed study of mineral chemistry in this study provides a new framework for understanding the

evolution of the Somerset Dam Igneous Complex. The collection of mineral chemistry placed into a

stratigraphic context allows for the identification and interpretation of variations in major, minor and

trace element chemistry thereby adding to the geological knowledge of the SDIC. Specifically, this

data set can address the existing hypothesis that the dominant mechanism driving the layering in the

SDIC was magma recharge. This was able to be achieved due to the unaltered state of the rocks, and

specifically the absence or minimal impact of subsolidus thermal re-equilibration.

The in situ data taken from the SDIC provides strong evidence of magma recharge within a large

troctolite unit, approximately 80 meters in thickness. This is shown by:

1) Samples taken from within this unit between 270m and 300m are often reversely zoned and

experience a sharp and contrasting increase in An component from ~An60 below the recharge

horizon to ~An80 above it.

2) Minor and trace element data show trends that are most easily explained by the interaction of

multiple liquids. Ba, Sr and Fe behave anomalously in the recharge zone (between 270m and

300m) which is hypothesised to be the result chemical buffering. Additionally REE patterns

from core to rim cannot be explained by simple fractional crystallisation.

3) Rock textures within the troctolite unit also support the hypothesis of magma recharge as the

samples in the recharge zone become strongly flow aligned and are evidence that fluid

transport was localised in this stratigraphic horizon and absent elsewhere. Additionally, this

section of the stratigraphy is characterised by reabsorption textures in the cores of plagioclase.

Because the signals of magma recharge can be found within the large troctolite unit instead of at the

contacts between macrocycles as is proposed by previous authors, a revised interpretation of the SDIC

is required.

1) We propose that the SDIC is the product of the amalgamation of sills into a ‘composite sill’.

Instead of each macrocycle being the product of a recharge event within a single magma

64

chamber, the evidence collected in this study fits an interpretation that each macrocycle is

instead a discrete sill.

2) These sills are intruded adjacent to each other and may have sharp differences in chemical

characteristics.

3) There is evidence that macrocycles 3 and 4 may instead be a single sill which experienced

magma recharge, with the new pulse of liquid intraplating the existing sill. This idea is

supported by the evidence demonstrating magma recharge in the central troctolite, but also by

the distribution of trace element concentrations in the stratigraphy which indicates that the

rocks sampled below the central troctolite crystallised from a different liquid than the rocks

within and above the central troctolite, and is potentially a different sill.

65

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Annen C (2005) The Genesis of Intermediate and Silicic Magmas in Deep Crustal Hot Zones. J Petrol 47:505–539. doi: 10.1093/petrology/egi084

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