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TRANSCRIPT
Monazite control on Th, U and REE
redistribution during partial melting:
experiment and application to the deeply
subducted crust
Aleksandr S. Stepanov
February 2012
A thesis submitted for the degree of Doctor of Philosophy
of the Australian National University
Declaration
The work presented in this thesis was carried out during my time in the Aus-
tralian National University. I certify that this thesis is my own work except
where otherwise acknowledged. Some of the ideas presented have benefited from
discussions with my supervisors and other colleagues, but all interpretations and
conclusions are my own.
Aleksandr S. Stepanov
Dedicated to my family and science
i
ii
Acknowledgements
I would firstly like to thank my supervisors, Daniela Rubatto and Joerg Hermann.
They offered me a fantastic project, provided background idea, methodology and
access to world class research facilities. At same time I enjoyed real freedom in my
research activities and received guidance only when I needed it. Daniela and Joerg
were very fast in making perfect corrections to my manuscripts and discussions
with them greatly enriched my understanding of the science. Andrey Korsakov
from the Institute of Geology and Mineralogy, Novosibirsk, Russia allowed access
to his comprehensive collection of Kokchetav samples, and generously provided
for this study the best samples ever. Andrey was always available to share his
knowledge regarding Kokchetav and UHP petrology. Our numerous discussions
greatly contributed to this work.
Bob Rapp is a man with unique insight into this project, as around 25 years
ago he published two papers on monazite solubility inherently relevant to this
study. Bob’s assistance was also crucial in achieving such accurate REE mea-
surements by electron microprobe. The advice I received from Ian Williams,
Trevor Ireland, Hugh O’Neill and Ian Campbell regarding experimental tech-
niques and monazite mineralogy is greatly appreciated. Bill Hibberson, Dean
Scott and Dave Clark provided technical support and assistance in performing
piston cylinder experiments. Under their comprehensive instruction I was able
to advance from a rather miserable 50% success rate, to 100% at the close of my
experimental program. The Centre for Advanced Microscopy at the Australian
National University, and Frank Brink in particular, is acknowledged for support
and access to SEM facilities. Charlotte Allen provided invaluable instruction and
support regarding LA-ICP-MS analysis. Terry Mernagh from Geoscience Aus-
tralia assisted with Raman spectroscopy. My thanks also go to Harri Kokkonen
and Shane Paxton for access to their lab facilities and for their sound advice on
how to handle the most problematic samples.
The help provided by David H. Green on my arrival in Canberra is something
I shall never forget.
The support of my fellow students at RSES is something I have greatly ap-
preciated, in their assistance with all aspects of my research project and in their
friendship during my time in Australia. Jesse Jones, Seann McKibbin, Evan
Gowan, Paolo Sossi, Paul Millsteed, Kate Kiseeva, Jason Doull, Kate Boston,
Jeremy Wykes, Tanya Ewing and Alex Mccoy-West contributed to the some-
what difficult editorial task of revising drafts of my manuscript. Don Dingwell
iii
and an anonymous reviewer for the journal “Chemical Geology” are gratefully
acknowledged for the editorial handling of what is the first chapter of this work.
I also wish to acknowledge the lecturers from my Alma Mater, the Novosi-
birsk State University, located in Novosibirsk Akademgorodok, Central Siberia.
Natalia Artemova Kulik, Olga Turkina, Vladimir Vasilievich Khlestov, Genadiy
Grigorievich Lepezin and Sergei Kargopolov introduced to me both the funda-
mentals and unresolved issues of petrology. Their lessons were truly modern,
progressive and informative. Thanks to their influence, I had little difficulty
in comprehending the new areas of research I was exposed to while attending
seminars or throughout the course of my own project.
Finally I wish to thank my family — Irina and Alisa. They are the beauty
and joy of my life. Irina and I have always been able to count on the support of
our parents in pursuing a scientific career, and I am deeply grateful to them. Our
parents also played a role in the production of this thesis through the crucial task
of babysitting Alisa on many occasions. HuiJuan, Magda, Juan Pablo, Tsuyoshi,
Gleen and the many other I became acquainted with during my residence in
Australia, you made my time in Canberra both enjoyable and memorable.
iv
Abstract
Rare earth elements (REE), Th and U are elements with similar geochemical
properties. In crustal rocks these elements are hosted by the light REE (LREE)
phosphate mineral monazite. Piston-cylinder experiments were conducted to
constrain monazite solubility and monazite/melt partitioning in hydrous granitic
melts at conditions relevant to anatexis at crustal conditions and in subduction
zones. Monazite has strong preference for LREE and Th; REE heavier than
Nd have decreasing compatibility in monazite, and U is less compatible than
LREE and Th. New experimental data and reconciliation with previous studies
led to a new formulation of LREE solubility in granitic melts as a function of
temperature, pressure, monazite composition and water content in melt.
The behaviour of monazite during high-pressure and ultra-high pressure (UHP)
metamorphism was studied using a suite of rocks from the Kokchetav massif,
Kazakhstan. Detailed petrographic and geochronologic study of the samples
from the UHP Kokchetav complex was combined with investigation of trace ele-
ment geochemistry and mineral inclusions of garnet, monazite and zircon. These
data demonstrated that (a) on prograde evolution rocks did not experience a
linear increase of pressure and temperature (P and T), but had stages of almost
isothermal increase of pressure and heating stage with a small increase of pres-
sure, (b) exhumation produced a close association of UHP gneisses with rocks
that experienced metamorphism at lower PT conditions and/or along different
paths from typical UHP rocks.
The geochemistry of the UHP gneisses of the Kokchetav complex is a perfect
target for the application of the new experimental data, because these rocks
experienced metamorphism and melting at the highest PT conditions recorded
in crustal rocks. Bulk rock geochemistry of the UHP gneisses shows pronounced
depletion in LREE, Th and U, and a smaller degree of depletion or enrichment
in other elements that are often considered as incompatible. The variation in
composition of UHP gneisses is explained by a new petrological model, which
takes into account the fact that restites are composed of residual assemblage
together with a residual melt. It is demonstrated that together with the residual
mineral association, the degree of melting and melt extraction efficiency play an
v
vi
important role in controlling of trace element behaviour.
Polyphase inclusions trapped in garnet were found in samples of some UHP
gneisses. The original composition of inclusions was obtained by high pressure
rehomogenisation experiments. The experiments demonstrated that polyphase
inclusions represent former melts of variable compositions, varying from high
temperature high-LREE melts formed at peak conditions to low-LREE melts
formed during exhumation. These inclusions are the first natural examples of
melts formed by melting of sediments at subarc depth. This partial melting led
to the complete dissolution of monazite and a strong depletion of LREE, Th and
U in the UHP gneisses. Melt inclusions and bulk rock geochemistry provide evi-
dence for the release of high LREE melts from melting of crustal metasediments.
Partial melting is thus an important process changing the physical and chemical
properties of deeply subducted crustal rocks.
____________________________________________________________________
Foreword to the Internet edition ____________________________________________________________________
I would like to thank three reviewers for thorough and detailed assessment of this work. Corrections were made whether possible and number of uncorrected issues will be improved or changed in publications produced from this manuscript. The text is still based on NIST values from Pearce et al. (1997) and concentrations will be corrected for new NIST values in journal versions of the chapters. However to the best of my knowledge these issues do not affect validity of observations and conclusions presented in this work.
September 2012, Canberra
Contents
Abstract v
Introduction 1
0.0.1 Abbreviations and symbols . . . . . . . . . . . . . . . . . . 5
1 Experimental study of monazite/melt partitioning with implica-
tions for the REE, Th and U geochemistry of crustal rocks 7
1.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7
1.2 Experimental and analytical techniques . . . . . . . . . . . . . . . 9
1.2.1 Starting compositions . . . . . . . . . . . . . . . . . . . . . 9
1.2.2 Experimental methods . . . . . . . . . . . . . . . . . . . . 10
1.2.3 Analytical methods . . . . . . . . . . . . . . . . . . . . . . 10
1.2.4 Experimental challenges and estimation of analytical errors 14
1.3 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15
1.3.1 Phase relationships . . . . . . . . . . . . . . . . . . . . . . 15
1.3.2 Monazite solubility in melts . . . . . . . . . . . . . . . . . 18
1.3.3 Monazite composition . . . . . . . . . . . . . . . . . . . . 19
1.3.4 Monazite/melt partitioning . . . . . . . . . . . . . . . . . 20
1.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25
1.4.1 Assessment of data quality and approach to equilibrium . . 25
1.4.2 Phase relationships . . . . . . . . . . . . . . . . . . . . . . 28
1.4.3 Comparison with previous experimental studies of mon-
azite solubility . . . . . . . . . . . . . . . . . . . . . . . . . 29
1.4.4 The general form of monazite solubility equation . . . . . . 29
1.4.5 New expression for monazite solubility . . . . . . . . . . . 35
1.4.6 Monazite/melt partitioning . . . . . . . . . . . . . . . . . 37
1.5 Implications of the experimental results . . . . . . . . . . . . . . . 41
1.5.1 Implications for fractional crystallisation of granites . . . . 41
1.5.2 Implications for crustal melting . . . . . . . . . . . . . . . 43
1.5.3 Implications for melting in subduction zones . . . . . . . . 45
1.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48
vii
viii
2 Association of rocks with different PT paths within the Barchi
Kol’ UHP block 59
2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59
2.2 Geologic setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60
2.2.1 Geology of the Kokchetav complex . . . . . . . . . . . . . 60
2.2.2 Geology of Barchi Kol’ unit . . . . . . . . . . . . . . . . . 63
2.3 Analytical methods . . . . . . . . . . . . . . . . . . . . . . . . . . 65
2.4 Strategy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66
2.5 Sample descriptions . . . . . . . . . . . . . . . . . . . . . . . . . . 67
2.6 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70
2.6.1 Mineral compositions and zoning . . . . . . . . . . . . . . 70
2.6.2 Mineral inclusions . . . . . . . . . . . . . . . . . . . . . . . 75
2.6.3 Monazite and zircon description and geochronology results 77
2.7 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88
2.7.1 Mineral inclusions in the Kokchetav rocks . . . . . . . . . 89
2.7.2 Application of Ti-in-zircon thermometers to UHP rocks . . 90
2.7.3 Interpretation of garnet zoning . . . . . . . . . . . . . . . 91
2.7.4 PT paths . . . . . . . . . . . . . . . . . . . . . . . . . . . 93
2.7.5 Linking accessory minerals with metamorphic evolution . . 99
2.7.6 Origin of rocks with different PT path inside UHP terrain 104
2.7.7 Tectonic implications . . . . . . . . . . . . . . . . . . . . . 106
2.7.8 Prograde PT path . . . . . . . . . . . . . . . . . . . . . . 110
2.7.9 Parallels with other UHP complexes . . . . . . . . . . . . 111
2.8 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114
2.9 Supplementary materials . . . . . . . . . . . . . . . . . . . . . . . 115
3 The geochemistry of Ultra High Pressure anatexis 117
3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117
3.2 Geological background . . . . . . . . . . . . . . . . . . . . . . . . 118
3.3 Analytical methods . . . . . . . . . . . . . . . . . . . . . . . . . . 119
3.4 Sample description . . . . . . . . . . . . . . . . . . . . . . . . . . 121
3.4.1 Non-UHP samples . . . . . . . . . . . . . . . . . . . . . . 121
3.4.2 UHP samples . . . . . . . . . . . . . . . . . . . . . . . . . 124
3.5 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125
3.5.1 Classification of non-UHP Kokchetav rocks . . . . . . . . . 125
3.5.2 REE minerals in UHP gneisses . . . . . . . . . . . . . . . 127
ix
3.5.3 Compositions of UHP gneisses and comparison with poten-
tial protolithes . . . . . . . . . . . . . . . . . . . . . . . . 128
3.6 Discussion: General . . . . . . . . . . . . . . . . . . . . . . . . . . 136
3.6.1 Host minerals for REE, Th, U in UHP rocks . . . . . . . . 136
3.6.2 Geochemical features of the Kokchetav sediments . . . . . 137
3.6.3 Constraining protolith composition . . . . . . . . . . . . . 139
3.7 Discussion: Behaviour of elements during UHP melting . . . . . . 141
3.7.1 Major elements . . . . . . . . . . . . . . . . . . . . . . . . 141
3.7.2 REE, Th and U . . . . . . . . . . . . . . . . . . . . . . . . 141
3.7.3 Large Ion Lithophile Elements: Rb, Cs, Sr and Ba . . . . . 143
3.7.4 High Field Strength Elements: Zr, Hf Nb and Ta . . . . . 145
3.7.5 Phosphorus . . . . . . . . . . . . . . . . . . . . . . . . . . 147
3.7.6 Sulfur, chalcophile elements . . . . . . . . . . . . . . . . . 147
3.7.7 Be . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147
3.8 Discussion: Numerical estimates on melting of UHP gneisses . . . 148
3.8.1 Melting model . . . . . . . . . . . . . . . . . . . . . . . . . 148
3.8.2 Estimate of melt loss . . . . . . . . . . . . . . . . . . . . . 151
3.8.3 Efficiency of melt extraction . . . . . . . . . . . . . . . . . 152
3.8.4 Modal abundances of minerals . . . . . . . . . . . . . . . . 153
3.8.5 Modal abundances of minerals and residual melt . . . . . . 153
3.8.6 Role of melting parameters on the behaviour of elements . 155
3.8.7 Melting history of the UHP gneisses . . . . . . . . . . . . . 156
3.8.8 Nomenclature of the Kokchetav migmatites . . . . . . . . 158
3.8.9 Comparison of Kokchetav UHP gneisses with other restites 160
3.9 Implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 162
3.9.1 Implications for the exhumation of the Kokchetav UHP
complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . 162
3.9.2 Implications for the trace element signature of subduction
zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 162
3.9.3 Implications for the isotopic signature of enriched mantle
reservoirs . . . . . . . . . . . . . . . . . . . . . . . . . . . 164
3.10 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166
4 Polyphase inclusions in an UHP gneiss 171
4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171
4.2 Analytical methods . . . . . . . . . . . . . . . . . . . . . . . . . . 172
4.3 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173
x
4.3.1 Sample description . . . . . . . . . . . . . . . . . . . . . . 173
4.3.2 Mineral compositions . . . . . . . . . . . . . . . . . . . . . 174
4.3.3 Polyphase inclusions . . . . . . . . . . . . . . . . . . . . . 181
4.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 191
4.4.1 Origin of polyphase inclusions . . . . . . . . . . . . . . . . 191
4.4.2 Interpretation of homogenization experiments . . . . . . . 193
4.4.3 Estimation of temperatures of inclusions formation . . . . 193
4.4.4 LREE, Th and U evolution during melting . . . . . . . . . 194
4.4.5 Hosts for LILE and their behaviour during melting . . . . 196
4.4.6 Hosts for HFSE and their behaviour during melting . . . . 197
4.4.7 Hosts for Be and its behaviour during melting . . . . . . . 199
4.4.8 Numerical estimates of melt loss . . . . . . . . . . . . . . . 199
4.4.9 Comparison with other data on fluid/melt inclusions from
other UHP rocks . . . . . . . . . . . . . . . . . . . . . . . 202
4.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 206
5 Conclusions 215
Bibliography 219
Chapter 1
Experimental study of
monazite/melt partitioning with
implications for the REE, Th
and U geochemistry of crustal
rocks
1.1 Introduction
Monazite is a Light Rare Earth Elements (LREE) phosphate (LREE)PO4 that
typically occurs as an accessory mineral in metapelites, granulites, peralumu-
ous granites and is also common in carbonatites and kimberlites (Lyakhovich &
Barinskii, 1961; Overstreet, 1967; Jones et al., 1996; Bea, 1996; Spear & Pyle,
2010). Monazite is the major host for LREE, Th and U in low-Ca granites and
metapelites (Overstreet, 1967; Bea, 1996). Monazite is also an important mineral
for Th, U–Pb geochronology of crustal rocks. It has recently been proposed that
monazite and/or allanite are important hosts for LREE, Th and U in deeply
subducted crustal rocks (Plank, 2005; Klimm et al., 2008; Hermann & Rubatto,
2009; Plank et al., 2009; Skora & Blundy, 2010).
Synthetic REE(PO4) compounds crystallize in the monoclinic monazite struc-
ture for rare earths from La to Gd, in which REE ions are located in 9-coordinated
polyhedra (Ni et al., 1995). Rare earths from Tb to Lu form crystals with a tetrag-
onal structure that is isostructural with xenotime (Y PO4) and zircon (ZrSiO4),
in which ions reside in smaller 8-coordinated polyhedra (Ni et al., 1995). HREE
and Y are the major impurities in monazite, representing the xenotime compo-
nent. An immiscibility gap exists between monazite and xenotime that shrinks
with temperature (Gratz & Heinrich, 1997). Pressure as well as the presence of
7
8
Th increases the solubility of Y in the monazite structure (Gratz & Heinrich,
1997; Seydoux-Guillaume et al., 2002). Thorium usually comprises 1–10 wt.%
of natural monazites, forming two different substitution mechanisms with en-
members huttonite (ThSiO4) and cheralite (Ca0.5Th0.5PO4). The endmember
ThSiO4 has two polymorphic modifications: thorite and huttonite. Tetragonal
thorite is isostructural with zircon and is stable at high T and low P; huttonite
is isostructural with monazite and is stable at high P and lower T (Taylor &
Ewing, 1978). However, huttonite can crystallize in the stability field of thorite
by epitaxial growth on monazite (Harlov et al., 2007). Monazite has low solubil-
ity in aqueous fluids, but it can increase in solutions containing NaCl (Ayers &
Watson, 1991; Poitrasson et al., 2004; Tropper et al., 2011).
The behaviour of accessory minerals in metamorphic and magmatic processes
can be described by three “fundamental accessory phase parameters”: solubili-
ties, diffusivities and the mineral/liquid partition coefficients (Watson & Harri-
son, 1984). Monazite solubility in granitic melts at crustal conditions has been
the subject of a number of experimental studies (Rapp & Watson, 1986; Rapp
et al., 1987; Montel, 1986, 1993). The diffusivities of LREE in granitic melts dur-
ing monazite dissolution were determined by Rapp & Watson (1986). The last
fundamental parameter, monazite/melt partitioning, has not yet been properly
considered experimentally for a wide range of trace elements, with only partition-
ing calculations from natural samples available (Yurimoto et al., 1990). Recently,
Skora & Blundy (2010) reported monazite/melt partition coefficients for LREE,
Th and U, over a range of temperatures, but at constant pressure (30 kbar).
Experimental studies of accessory minerals are complicated by a tendency to
grow small grains, their complex composition, and by very slow solid-state diffu-
sion (Watson & Harrison, 1984; Rubatto & Hermann, 2007a; Tailby et al., 2011).
In nature, accessory minerals often show inheritance and disequilibrium features.
Fine scale zoning and disequilibrium processes at the crystal-melt interface have
the potential to bias the results as well. However, recent advances in high reso-
lution analytical techniques allow these issues to be adequately addressed.
In this chapter, I present the first experimental study to examine monazite/melt
partitioning and monazite solubility over a wide range of P-T conditions. Piston-
cylinder experiments were performed over a temperature range from 750 to
1200oC and a pressure range from 10 to 50 kbar. Experimental charges were
analyzed by Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LA-
ICP-MS) and by Electron Probe Micro Analysis (EPMA); where direct measure-
ment of monazite composition was impossible due to its small crystal size, its
9
composition was calculated by regression analysis of the LA-ICP-MS data. From
these analyses, I have determined the solubility of monazite in hydrous granitic
melts as a function of temperature and pressure, and monazite/melt partition
coefficients for a wide range of trace elements. These data are briefly applied to
granite crystallisation, and to melting at crustal conditions and in subduction
zones.
1.2 Experimental and analytical techniques
1.2.1 Starting compositions
The starting material was produced using a ’sol-gel’ method to eliminate prob-
lems associated with the sluggish reaction kinetics of refractory minerals during
experiments. Most major elements and trace elements were combined as nitrate
solutions, then mixed with tetraethyl orthosilicate [Si(C2H5O)4] and slowly dried
to a gel. Since melt composition has been shown to play an important role in
monazite solubility (Rapp et al., 1987; Montel, 1986), a peraluminous granitic
composition free of Fe and Mg has been chosen as the starting material, so that
the results are particularly relevant to crustal anatexis and granite formation
(Table 1.1). LREE, Th and U were added to the mix in ratios similar to the
composition of natural monazite, whereas HREE were doped to a higher level
than in natural monazite in order to increase concentrations in both melt and
monazite to levels more easily measured by EPMA and laser-ablation ICP-MS.
The mix was then ground in an agate mortar and melted to a glass in a platinum
crusible at 1400oC. Aluminum was then added to the ground glass as Al(OH)3
or Al2O3, in order to accurately control the amount of water in each experiment.
Identical compositions were then made without trace elements.
The reconnaissance experiments have shown that if the content of trace ele-
ments in the starting material is too high, then dispersion of tiny monazite grains
complicates the measurement of ”pure” glass composition by microbeam meth-
ods. By mixing variable amounts of the different starting materials in Table 1.1,
I was able to control the amount of monazite (0.5-3.0 wt.% monazite) and water
in the system, without changing the overall bulk composition of the melt. Sev-
eral additional experiments were performed in water-rich bulk compositions. For
these experiments, starting compositions were prepared from anhydrous compo-
nents and distilled water was added then with a microsyringe.
10
1.2.2 Experimental methods
Experiments were performed in an end-loaded piston-cylinder apparatus at the
Research School of Earth Sciences, Australian National University (ANU). Ex-
periments at 10–40 kbar were conducted in standard 200 T hydraulic presses,
whereas the 50 kbar experiments employed a 500 T, ultrahigh-pressure press.
For all experiments, 15–20 mg of sample was placed in a � 2.3 mm Pt capsule
that was then sealed by arc welding. Sample capsules in most cases were ≈6
mm long, except experiments with free water, which used ≈10 mm long cap-
sules. The capsule was kept cool (below 100oC) during welding by immersion
in water-soaked tissue paper. The capsule end was folded to produce an addi-
tional cold seal. The 1/2" or 5/8" cell assemblies consisted of teflon film, NaCl
with or without Pyrex sleeves (dependent upon whether the experiment was per-
formed above the NaCl melting curve or below), a graphite heater and MgO
spacers above and below the capsule. The Pt capsule was placed inside a cylin-
drical MgO sleeve and the void space filled with MgO powder. Care was taken
to position the capsules within the hotspot of the assembly, within ≈0.5mm of
the thermocouple tip. Temperature was monitored with type-B thermocouples
(Pt94Rh6/Pt70Rh30) and regulated using a Eurotherm temperature controller
considered to be accurate within ±3oC. Pressure was converted directly from
the load and was accurate to ±1 kbar. Pressure was adjusted several times dur-
ing the first 24 h of each experiment to compensate for any loss of pressure due to
the relaxation of internal friction. Experiments with Pyrex were initially heated
to 600oC under low confining pressure (5 kbar) in order to soften the Pyrex,
before temperature and pressure were increased simultaneously up to the desired
run conditions.
1.2.3 Analytical methods
The experimental samples were set in epoxy, sectioned, and polished prior to anal-
yses using the JEOL 6400 scanning electron microscope (SEM) at the Centre for
Advanced Microscopy, ANU. Back-scattered electron images of the samples were
obtained, and the major-element composition of all phases were determined using
an accelerating voltage of 15 kV and a beam current of 1 nA. The composition
of glasses was determined using area scans > 5× 5 μm in order to avoid loss of
Na and K during the analysis. The H2O content of the glasses was estimated
from the deviation of the oxide totals of the melt from 100% (Devine et al.,
1995). Experiments with high water content and with significant crystallization
11
produced vesicular glass on the quench. In these charges, SEM analyses by by
areal scanning were performed as much as possible over relatively bubble-free
areas of quenched glass. The water content of the melt in these experiments has
also been calculated based upon mass balance, and these estimates agree within
1.5% with the estimates obtained by the difference method.
The concentration of selected REE, Y, Th, U and P in glasses and monazite
were determined by wavelength dispersive spectrometry (WDS) using the Cameca
SX100 electron microprobe at the RSES. Beam current and accelerating voltage
were 100 nA and 15 kV, respectively, and the spot size was ≥ 3μm. Different
acquisition procedures were used for experimental monazite and glasses. X-ray
emission spectra of monazite was acquired and used for selection of the peak and
background positions for P, Y, La, Ce, Pr, Nd, Sm, Th and U. These elements
were measured in monazite using peak counting times of 10–60 s. For the glasses,
La and Ce concentrations were simultaneously measured on two LPET crystals
using counting times of 60–240 sec, which resulted in a detection limit of about
100 ppm Ce (Fig. 1.1). In most experiments, the concentration of LREE in
the glass was measured by both LA-ICP-MS and electron microprobe, and show
remarkably good agreement (Fig. 1.1a,b). In order to reveal monazite zoning,
exceptionally large crystals were BSE imaged with a high sensitivity Cambridge
S360 SEM operating at 20 kV acceleration voltage, a beam current of 2 nA and
a working distance of 15 mm.
Concentrations of Si, Ca, Al and trace elements of coexisting melt and mon-
azite in experimental charges were determined by LA-ICP-MS at the Research
School of Earth Sciences (ANU) using a pulsed 193 nm ArF Excimer laser with
100 mJ energy, operating at a repetition rate of 5 Hz (Eggins et al., 1998), cou-
pled to an Agilent 7500 quadrupole ICP-MS. Laser sampling was performed in a
He-Ar atmosphere using a spot size of 16–25 μm. Data acquisition was performed
by peak hopping in pulse counting mode, acquiring individual intensity data for
each element during each mass spectrometer sweep. For each analysis data were
collected over total time of 60 sec., including gas background of 20–25 sec.
BSE images were used as a guide to search for areas of clean glass or areas rich
in monazite crystals during LA-ICP-MS analysis. Up to 20–25 individual spot
analysis were made on each experiment. In a count rate versus time diagram,
segments of 5–25 second duration were integrated. Synthetic glass NIST 610
was used as the external standard material, using values taken from Pearce et al.
(1997). The SiO2 and CaO contents of the glass, as measured by EDS, were used
as internal standards. The standardisation using SiO2 showed smaller variation
12
Figure 1.1: Comparison of trace element analyses in glasses (a,b) and in monazite(c) performed by LA-ICP-MS and by WDS EPMA. Error bars are 1 standarddeviation.
in the trace element content on crystal-free glass and was adopted as the internal
standard for the mixed analyses.
In most experiments, direct measurement of individual monazite crystals by
LA-ICP-MS trace element compositions was impossible because the width of the
monazite crystals was less than the diameter of the laser beam. This problem was
overcome by analyzing monazite-melt mixes by LA-ICP-MS (Fig. 1.2) followed
by statistical treatment of mixed analyses. Monazite compositions from mixed
mineral-melt LA-ICP-MS analyses were determined using the regression method
described by Rubatto & Hermann (2007a). The entire dataset was linearly re-
gressed relative to Ce (Fig. 1.2), with regression lines extrapolated to the Ce
content of monazite determined by EPMA (25.8 wt.% Ce), and the concentra-
tions of individual trace elements calculated on the basis of that extrapolation.
The error introduced assuming a single, constant Ce content for all the exper-
imental monazites is discussed below (section 1.4.1). All the regression trends,
13
Figure 1.2: Regressions between Ce and U (a,b) or Lu (c–f) for mixed melt-monazite analyses. Figures on the right represent enlarged parts of the diagramsfrom the left. Figures a-d show experiments performed at different pressureat temperature of 1000 oC and figures e and f show experiments at differenttemperatures at pressure of 10 kbar.
which are based on the LA-ICP-MS analyses only and do not include the end-
member monazite, were checked on binary diagrams and in general have r2 values
> 0.95; in the worst cases, r2 remains > 0.75.
An alternative to the regression analysis method employed above is the ”sub-
traction method”. From the mixed analysis, the melt composition is subtracted
14
and projected to monazite composition using the Ce content of monazite as an
internal standard, and the average calculated. This method works well for LA-
ICP-MS analyses in which the amount of crystalline monazite being ablated
(along with glass) is high; however, this approach is subject to large propogation
errors when the amount of monazite is low.
1.2.4 Experimental challenges and estimation of analyti-
cal errors
Calculation of monazite composition from monazite melt mixes has internal lim-
itations. Estimates can only be achieved for elements significantly enriched in
monazite; the lower the enrichment in monazite, the higher the error. In practice,
the lowest monazite/melt partition coefficient obtained by regression was around
10. For all the experiments, meaningful regressions were obtained for the REE,
Th, U and Y. Arsenic content in monazite was calculated for several experiments,
however its concentration in the melt was below the detection limit because of
high background and low concentration. For other trace elements, concentration
in monazite appeared below the detection limit of the regression method (see
below discussion of notable exceptions).
LA analysis of glasses in the low temperature experiments was complicated by
a dispersion of tiny monazite crystals and the presence of other silicate minerals.
Because monazite is so extremely enriched in trace elements, contamination of the
glass analyses by even a few of these small crystallites will produce a large eleva-
tion in the apparent concentration of REE measured in the glass. This problem is
analogous to the micronugget effect that complicates many experimental studies
of trace element partitioning (O’Neill et al., 2008). Due to this effect, there was
considerable dispersion in the concentration data at the low end of the composi-
tional range studied, with a continuous transition from analysis spots with low
LREE concentrations (attributable to dilution of the analysis by silicate phases),
to high REE concentrations (caused by contamination of the analysis by small
amounts of monazite crystallites). In some cases, the LA datasets did not provide
a clear indication as to which spots should be used to calculate the composition of
the glass. To avoid bias, I used La and Ce concentrations in the glass determined
independently by electron microprobe, which has greater spatial resolution. The
full set of low-LREE analyses were treated by the regression calculations, and the
concentrations of the other trace elements in the melt were interpolated based
on the Ce content, as determined by electron microprobe.
15
Some experiments crystallized epidote group minerals, apatite, and zircon
(Table 1.2). As these minerals contain significant amounts of REE, their pres-
ence in glass-monazite mixes can affect calculations of the monazite composition.
However, in all experiments monazite is more abundant than the other accessory
phases, and apatite and epidote group minerals usually were found near the bot-
tom of the capsule. Moreover, in all experiments, regressions between Ca, P, Zr
and LREE were also checked, and those analyses deviating from the major trend
between monazite and melt were excluded.
Error propagation for monazite composition was calculated using the formu-
lae ”Prediction Intervals for Individual Estimates” from Helsel & Hirsch (2002)
at 95% confidence level. In a number of experiments, the propagated error on
monazite composition is as low as 1.5%. This is surprising, because the ex-
pected precision of LA-ICP-MS is 5-10% and the large extrapolation can only
diminish this. However, regression is essentially based on element ratios, which
LA-ICP-MS can measure with greater precision than absolute concentrations.
Experiments conducted at lower temperatures or with low water contents pro-
duced very small monazite crystals, resulting in a lower maximum content of
monazite in the mixed analyses, and a smaller range of compositions, so that
extrapolating to monazite compositions leads to relatively large error.
In a few experiments it was possible to analyse monazite by EPMA without
any contamination from the melt. These analysis show good agreement with
the LA data. In particular, the Th content of monazite calculated by regression
is apparently invariable, and EPMA gives similar average Th concentration in
monazite. Also there is agreement in the HREE and Y contents of monazite
measured by EPMA, and by regression of the LA-ICP-MS data (Fig. 1.1c).
1.3 Results
1.3.1 Phase relationships
Most experiments produced small euhedral grains of monazite unevenly dis-
tributed in the glass (Fig. 1.3a–c). The size of these monazite crystallites
decreases with decreasing temperature and water content in the melt. In ex-
periments below 1000oC, crystals were less than 2 μm across (Fig. 1.3a–b), and
in experiments above 1000oC, they were up to 5 μm. A notable exception was
experiment C3970 at 1000oC, 10 kbar and 17 wt.% water, which produced excep-
tionally large monazite crystals (up to 100μm; see Fig. 1.3d). Monazite crystals
16
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50 μm50 μm
50 μm 20 μm
Figure 1.3: SE images of experimental runs: (a) monazite (bright crystals) inglass with plagioclase and quartz. q = quartz, pl = plagioclase, m = melt;(b) (800 oC, 10 kbar, 16% H2O) monazite (bright crystals) in glass containingexsolution bubbles; (c) (1000 oC, 50 kbar) central part of capsule with melt,bright crystals = monazite and ep-min = epidote group mineral; (d) BSE imageof large monazite crystals that show weak sector zoning.
grew close to the top and bottom of the capsule but are absent in the centre.
Crystallisation of large monazite in this experiment can be explained by the com-
bination of two factors: reduced viscosity, which speeds up diffusion, and a small
temperature gradient in the long capsule resulted in transportation of monazite
components to the tips of the capsule.
Apart from monazite and melt, which are present in all experiments, other
phases crystallised in several runs (Fig. 1.3, Table 1.2). Quartz and plagioclase
were present in experiments C3308 and C3322 (10 kbar, 750 and 800oC with
10 wt.% of H2O) and in experiment C3984 (10 kbar, 1000oC with 4 wt.% of
H2O in the system). Feldspar grains in these experiments are not homogeneous
and often contain inclusions of Al2O3 aggregates. These aggregates probably
17
represent unreacted residue from the dehydration of the Al(OH)3 that was added
to the original sol-gel starting material as a water source.
In the higher pressure experiments (P > 20 kbar), significant amounts of
silicate phases crystallised but the melt fraction always was above 50%. Silicate
crystals were usually concentrated at the top and the bottom of the capsule,
and a large pool of melt with monazite was generally observed in the centre.
At 30 kbar, coesite and kyanite start to crystallize at 1000oC, and at 900oC an
epidote group mineral is present (Table 1.3). In run C4028 (30 kbar, 750oC, 16
wt.% H2O), muscovite is present (see Table 1.3). Coesite appears at 1200oC,
50 kbar, and the assemblage coesite, kyanite and jadeite is observed at 1000oC.
In most experiments above 20 kbar and temperatures in the range 750–1000oC,
epidote group minerals with formulae close to Ca2Al3Si3O12(OH) (Table 1.3) are
present. Epidote group minerals from different experiments contain 1–2 wt.% Ce
and 0.7–1.3 wt.% La, as determined by EPMA. This mineral should be classified
as clino-zoisite–(Ce). The LREE content of the epidote group minerals is rather
similar in different runs, with no systematic variation in LREE content with
respect to T or P.
Saturation of the melt in zircon was reached in experiments at 800 and 750oC,
both at 30 and 10 kbar, with 60–100 ppm Zr in the melt. At pressure above 20
kbar, crystallisation of silicates concentrated incompatible elements in the melt.
In experiment UHPPC-142 (1000oC, 50 kbar), Zr concentration reached 370 ppm,
although zircon was not observed. Apatite was found in experiments UHPPC-142
and C3337 (1000oC, 50 and 40 kbar).
In all experiments, the melt has a major-element composition that is broadly
granitic, with some variation due to crystallisation of silicate phases. At 10 kbar,
crystallisation of quartz and feldspar produces melt with higher water content but
constant ASI (Alumina saturation index: molar ratio Al/(Na+K+2Ca)) (Table
1.2). At 30 kbar, the liquidus was crossed between 1100 and 1000oC and crys-
tallisation caused an increase in water and alkali elements, with an associated
decrease in alumina, resulting in lower ASI indexes (0.92-0.98) and melts that
are metaluminous. Increasing pressure to 50 kbar reduces the alumina content
even further, resulting in melts with ASI of 0.85. Glasses in experiments con-
ducted at low temperatures or with high water contents contain small bubbles.
Small size of bubbles and their homogeneous distribution across capsule indicates
that they formed by exsolution of water from fluid-rich melt during quenching
(Fig. 1.3b).
18
Figure 1.4: (a) Arrhenius plot of Ce concentration in melts buffered by monazite.Data from experiments by Rapp & Watson (1986) are plotted for comparison.(b) Effect of pressure on monazite solubility demonstrated by experiments whereonly pressure was different and melt composition was identical.
1.3.2 Monazite solubility in melts
Data on monazite solubility in melts are provided in Table 1.4 and represented
graphically in figure 1.4. LREE and phosphorus contents in the melt show a
strong temperature dependence, with Ce content in the melt varying from 86 to
3,000 ppm. The ratio of P/Ce in the melt decreases with increasing temperature.
The effect of water on the solubility of monazite in granitic melts is more
difficult to assess from these data. In the experiment with high water content
(C3970: 1000oC, 10 kbar, 17% H2O), the amount of Ce dissolved in the melt is
930±42 ppm, which is within the error of the amount of Ce in the melt when
there is only 10 wt.% H2O present (1050±53 ppm, C3309). No zoning in the
trace element compositions of the melts has been observed throughout the cap-
sules in these experiments. However, in the experiment with the lowest water
content (C3984: 1000oC, 10 kbar, 5.9% H2O), monazite solubility has decreased
significantly (to Ce 511±51 ppm). In experiments at 10 kbar and 800oC, an
increase of water content from 10 wt.% to 17 wt.% clearly enhanced monazite
solubility from 370 to 620 ppm.
The effect of pressure on monazite solubility (Fig. 1.4b) can be most directly
constrained from experiments at 1200-1000oC and 20–30 kbar that have identical
melt composition. Experiments at temperatures <1000oC and >30 kbar are less
suitable for this purpose because with increasing pressure, crystallisation of other
silicate phases shifts the melt to more hydrous and alkaline compositions. Both
of these factors can increase monazite solubility, masking the effect of pressure.
19
Comparison of the experiments performed at 10 and 30 kbar and temperatures
of 1100 and 1200oC show that increasing pressure from 10 to 30 kbar decreases
monazite solubility by 25–45%.
1.3.3 Monazite composition
Direct LA-ICP-MS analysis of large monazite crystals from experiment C3970
allowed us to determine a full list of trace elements in monazite. The data show
high concentrations of all REE, Th, U and Y, and also there is a strong positive
correlation between LREE, As and V. As and V concentrations in monazite
are 300–500 ppm and 56 ppm, respectively. Also Sr and Ti have measurable
concentrations in monazite (21 and 7 ppm respectively) but lower than in the
melt. Other trace elements from the starting composition (Li, Be, B, Sc, Mn,
Zr, Nb, Ba, Hf, Ta and Pb) have concentrations in monazite that are below the
limits of detection for the LA-ICP-MS, which suggests monazite/melt partition
coefficients less than 0.1.
In the experimental monazite, the main constituents (calculated as molar
fraction from the sum REE+Th+U+Y) are the LREE (e.g., La–Sm: 0.8–0.86),
the MREE (Eu–Dy: 0.026–0.04), the HREE (Ho–Yb: 0.0056–0.0124), and Y
(0.034–0.078). Th and U content is in the range of 0.058–0.079 and 0.0009–0.004,
respectively. Monazite contains≈1 wt.% SiO2 and 0.7–1.2 wt.% CaO. 40–70% of
the Th enters the experimental monazite as a huttonite component, and the rest
as a cheralite component. Over the range of conditions studied, monazite shows
significant variation in its HREE, Y and U content, which increase simultaneously
and are negatively correlated with LREE, while the concentration of Th remains
relatively constant.
In most of the experimental samples, monazite was too small for zoning to
be apparent; however, the exceptionally large crystals of monazite in experiment
C3970 (1000oC, 10 kbar, 16% H2O; Fig. 1.3d) do show zoning, allowing us to
evaluate the scale and compositional variability of these different zones directly.
BSE images of the large monazite crystals reveal sector zoning, with different
levels of brightness corresponding to the different monazite composition (the
brighter the area the higher the average atomic number). LA-ICP-MS analyses
of large monazites in C3970 provided pure monazite composition, with little or no
contamination by the melt. LA data show relatively large scatter, which is caused
by sector zoning of these crystals, but the laser’s spatial resolution is not sufficient
for analysis of separate zones. Microprobe analysis confirmed chemical zoning
20
in monazites. Variation of composition can be expressed as relative variation:
(Cmax − Cmin)/Caverage. For the LREE, the variation is relatively small: 4.5–
11.5%. However for Sm, Gd and Y, relative variation increases to 30–40%. Th
is also variable from 4.4 to 8.5 wt.% and is positively correlated with Si. In
general, LREE and HREE are negatively correlated. These correlations show
that in different growth sectors there are independent substitutions of LREE by
ThSiO4 and LREE by M-HREE.
1.3.4 Monazite/melt partitioning
The primary outcome of this study is the first experimental determination of
monazite/melt partition coefficients for the REE, Th, U and Y. The composi-
tion of monazite is presented in Table 1.5 and the composition of the co-existing
melts are given in Tables 1.2 and 1.4. Distribution coefficients between monazite
and melt are reported in Table 1.6. The values obtained cover large tempera-
ture (750–1200oC) and pressure (10–50 kbar) intervals. Partition coefficients in
different experiments are shifted relative to one another because the solubility
of monazite varies from experiment to experiment. It is convenient to compare
the monazite/melt partitioning values in different experiments by normalizing
against a single value for one element. In Figures 1.5, 1.6 and in the follow-
ing discussion, all partition coefficients are normalized to an arbitrary value of
Dmnz/lCe = 1000 (denoted as K
mnz/lD X/Ce(1000), after (Beattie et al., 1993)).
Generally, monazite/melt partitioning patterns are similar in all experiments.
LREE show a very limited range in partition coefficients (Fig. 1.5, 1.6), and
elements heavier than Nd have a progressively weaker preference for monazite.
Monazite strongly concentrates Th but has a lower preference for U. This simple
picture is complicated by some variations in the partitioning of specific elements
at different conditions.
It is important to determine if there are any differences in monazite/melt
partition coefficients among the LREEs. This is deduced by looking at element
ratios in melt and coexisting monazite (Fig. 1.7). For example, a plot of La/Nd
vs. Sm/Nd shows that there is no systematic difference in the La/Nd ratio of
melt and monazite. On the other hand, the Sm/Nd ratio of the melts are higher
than those of monazite, with the starting material in between. Therefore I can
conclude that for La, Ce, Pr and Nd, the difference in partitioning is less than
the precision in the most precise data, e. g. less than ≈5%. These properties of
monazite appear to be independent of T and P.
21
Figure 1.5: Partitioning of trace elements between monazite and melt. (a)
Dmnz/melt normalized to Dmnz/meltCe = 1000 ratios. (b) comparison of measured
Dmnz/melt with published values from Yurimoto et al. (1990), Ward et al. (1992),Bea et al. (1994) and TS – this study. KD(mnz/melt)* – monazite/melt parti-
tioning coefficients normalized to Dmnz/meltCe =1000
Samarium is apparently the first REE with partition coefficients measurably
lower than LREE. Kmnz/lD Sm/Ce(1000) ranges from 1000 to 590 under different condi-
tions. The weighted average of Kmnz/lD Sm/Ce(1000) is 800± 36 (mean square weighted
deviation — MSWD is 2).
For yttrium, the Kmnz/lD Y/Ce(1000) varies from 400 to 113. The systematic varia-
tion in the Y content of monazite in different experiments, confirmed by both LA-
ICP-MS and EPMA (Fig. 1.1c), is strong evidence for variation of Kmnz/lD Y/Ce(1000).
The majority of experiments have Kmnz/lD Y/Ce(1000) close to 200. The low temper-
ature experiments at high pressure have Kmnz/lD Y/Ce(1000)=110–140. K
mnz/lD Y/Ce(1000)
apparently increases with temperature and decreases with pressure, although
greater accuracy is necessary in order to confirm these trends. The weighted
average of Kmnz/lD Y/Ce(1000) is 206±25 with a MSWD of 24; this scatter cannot be
attributed to analytical uncertainty. With the exception of five experiments,
22
Figure 1.6: Effects of temperature and H2O contents on of monazite/melt traceelement partitioning. The values are monazite/melt partitioning normalized to
Dmnz/meltCe = 1000. Different symbols denote different pressures and rev is a
reversal experiment, performed with natural monazite at 1200 oC , 10 kbar.
23
Figure 1.6: Continued.
Kmnz/lD Y/Ce(1000) is within the narrow interval of 150–320. Yttrium has partition co-
efficients slightly lower than Ho, in agreement with the close ionic radii of these
elements (Shannon, 1976).
The HREE have lower concentrations, larger error and a greater disper-
sion of the data relative to Y, but they essentially show behaviour similar to
Y. Kmnz/lD Lu/Ce(1000) varies from 30 to 160, with a statistical scatter greater than
the analytical uncertainty. There is a weak positive correlation with T and a
negative correlation with the water content of the melt. Monazite in almost
all experiments has a strong negative europium anomaly relative to the starting
composition, indicating relatively reducing conditions in the experiments, with a
significant amount of Eu present as Eu+2 (Rubatto & Hermann, 2007a).
Partitioning of thorium in monazite is always stronger than for the LREE,
with Kmnz/lD Th/Ce(1000) 1100–1700. K
mnz/lD Th/Ce(1000) does not show any systematic vari-
24
Figure 1.7: Plots representing element ratios in melts and monazites relative tostarting composition. (a) La/Nd vs. Sm/Nd; the La/Nd ratio has the same rangeboth in melt and in monazite, and the Sm/Nd ratio is higher in melts, than inthe starting composition and lower in monazites. (b) Y/La vs. Th/La; Th/La ofmonazite is slightly higher than in melt and Y/La ratio is strongly fractionatedand varies between different experiments.
25
ation with pressure, temperature or melt composition. Low and high Kmnz/lD Th/Ce(1000)
values were measured in experiments with small monazite grains and relative large
uncertainties and thus are not very reliable. Monazite has higher Th/La ratios
than the coexisting melt, with the starting composition between the two (Fig.
1.7). Therefore it appears that Kmnz/lD Th/Ce(const) is constant in the experiments, and
the weighted average Kmnz/lD Th/Ce(1000) calculated for all experiments is 1294 ± 68
(MSWD 2.3).
Partitioning of uranium in monazite varies widely in the experiments, with
Kmnz/lD U/Ce(1000) ranging from 50 to 280. The K
mnz/lD U/Ce(1000) has a positive correla-
tion with T and a negative correlation with pressure. No correlation with the
water content of the melt was observed. At temperatures lower than 1000oC, the
Kmnz/lD U/Ce(1000) falls within a narrow range (90–130).
1.4 Discussion
1.4.1 Assessment of data quality and approach to equilib-
rium
The attainment of equilibrium in my experiments is evident from the similar size
of monazite crystals across the capsules and their euhedral shape. Large monazite
crystals from experiment C3970 display sector zoning and are compositionally
variable. This zoning is caused by differential uptake of trace elements by different
crystal faces (Cressey et al., 1999), and concentric and oscillatory zoning have
not been observed in the experimental monazites. Sector zoning in monazite
shows that its composition was controlled by equilibrium partitioning between the
crystal surfaces and melt and not by disequilibrium growth from a boundary layer
or by variable diffusivity of components to the crystal-melt interface (Watson &
Mueller, 2009).
Differential uptake of components by different facets could be a problem for
calculations based on crystal/melt partitioning, because the factors controlling it
are not clear. However, because my data were obtained by LA-ICP-MS, which
completely vaporizes one or more grains of monazite, intragranular variations in
composition are effectively averaged. In the case of experiment C3970, where
individual crystals of monazite were bigger than the diameter of the laser beam,
there are significant variations in the composition of different sectors, but the
average composition is close to that of monazite in other experiments (Fig. 1.8).
Therefore differential uptake of trace elements by facets should not be a signifi-
26
Figure 1.8: (a) Comparison of the partitioning calculated for monazite compo-sition measured directly by LA-ICP-MS and calculated by regression. (b) Com-parison of the monazite/melt partitioning in reversed and synthesis experiment.There is good agreement between reversed and synthetic experiments, and thecomposition of the starting natural monazite is significantly different from themonazite grown in the reversed experiment. KD(mnz/melt)* – monazite/melt
partitioning coefficients normalized to Dmnz/meltCe =1000
cant problem, because I have essentially determined partitioning between whole
monazite grains and the melt.
In order to assess the approach to equilibrium in the experiments, I performed
a ”true reversal” experiment with natural monazite. An appropriate amount of
natural monazite was added to a water-bearing granite without trace elements
(Table 1.1) and ground in an agate mortar for one hour. In order to ensure
saturation, grains of natural monazite also were added to the bottom of the cap-
sule. This experiment was run at 10 kbar and 1200oC for 1 week, and produced
melt containing euhedral monazite grains 5 μm in size and large monazite grains
with a corroded surface. New monazite most probably formed by a dissolution-
precipitation mechanism. LA-ICP-MS analysis of the reversal experiment shows
that the glass is homogeneous and the∑
LREE content in the melt (Ce 3880
ppm) is similar to the LREE content measured in corresponding synthesis ex-
periment (Ce 3040 ppm). LA-ICP-MS analysis of monazite in the reversal ex-
27
periment shows that there is significant inheritance from the starting monazite.
New monazite has lower HREE and U contents, similar LREE and slighter higher
Th compared to the starting monazite, and the inherited monazite has high Pb
content. Monazite/melt partition coefficients for the new monazite are similar
to the values obtained in the synthesis experiments (Table 1.6, Fig. 1.7, 1.8).
Therefore the reversal experiment provides strong evidence for the attainment of
equilibrium in my experiments.
The trace element fractionation caused by monazite crystallisation can be
assessed by plotting ratios of elements in monazites and glasses (Fig. 1.7). On
such a diagram, elements not fractionated by monazite should have the same ratio
in both monazite and glass, and fractionated elements will be distinguishable from
the starting composition. Importantly, La/Nd and Sm/Nd ratios in monazite
and melt are measured independently. Moreover, the fact that Sm/Nd ratios for
the melt and monazite are distinct strongly suggests that fractionation of these
elements is independent of propagated error.
The increase in error on monazite and melt compositions with decreasing
temperatures is a direct consequence of the minute grain size of monazite in
the low temperature experiments. However, in my opinion this increased error
does not imply that the patterns observed at high temperature (e. g. relative
fractionation of Sm-Nd and HREE, the absence of fractionation in the LREE)
change drastically at low temperature.
The percentage of Ce (molar CePREE+Th+U
) in experimental monazite is 36–
40%. An error of 10% is introduced by calculating the composition of monazite
assuming the same concentration of Ce, which is insignificant relative to other
sources of uncertainty.
Calculating experimental accessory mineral composition by regression analy-
sis is a relatively new approach; its validity has been previously assessed for zircon
by Rubatto & Hermann (2007a) and was shown to provide reliable results. Data
from this study show the same evidences for equilibrium as shown by Rubatto &
Hermann (2007a) for zircon, that is:
• (i) The composition of monazite calculated by regression shows good agree-
ment with microprobe analysis of monazite uncontaminated by glass com-
ponents.
• (ii) Monazite/melt partitioning values change smoothly from LREE to
HREE.
28
New data in this chapter provide three additional checks for the reliability
and consistency of the regression approach:
• (iii) Ratios such as Th/La and Sm/Nd independently determined for mon-
azite, melt and the starting material fall on a single line in accordance with
mass balance consideration (Fig. 1.7).
• (iv) the large monazite crystals in experiment C3970 are sector-zoned, not
concentrically-zoned.
• (v) Results from the true reversal experiment showed good agreement with
the those of the synthesis experiment.
1.4.2 Phase relationships
The results of this study can be considered representative of the phase relation-
ships of water-rich granites. The phases present at 10 kbar are quartz, feldspar
and zircon, and at ≥30 kbar the crystallisation sequence is coesite, kyanite, epi-
dote group mineral and muscovite/jadeite. For the starting composition with 10
wt.% H2O, the liquidus lies between 800 and 900oC at 10 kbar, is approximately
1000oC between 20 and 30 kbar, and is above 1200oC at 50 kbar. These results
are in broad agreement with the experimentally determined phase relationships
for S-type leucogranites at pressures up to 35 kbar (Huang & Wyllie, 1981). Some
differences (in particular, the order of appearance of kyanite and epidote group
minerals) are due to the higher calcium and aluminum content of the starting
material.
Although decreasing temperature at 10 kbar causes crystallisation of silicates,
the melt preserves its original peraluminous composition. At higher pressure,
crystallisation of coesite, kyanite and epidote group mineral causes the melt to
become more alkaline, and the ASI approaches unity. At 50 kbar and 1000oC,
the appearance of jadeite drives the melt towards compositions that are meta-
luminous, almost peralkaline, and strongly-enriched in potassium. This trend is
similar to that described for melts unbuffered by a K-bearing phase (Hermann
& Green, 2001), only more pronounced. My experiments are unbuffered with
respect to potassium, because the absence of Mg and Fe in the starting mate-
rial prevents formation of phengite in most experiments. The only exception is
experiment C2309 (750oC, 30 kbar), where a muscovite-like mica is present.
Apatite crystallisation can be explained by a rapid decrease in the concen-
tration of P in the melt relative to the LREE. Appearance of monazite in high
29
pressure experiments and allanite at lower pressures was observed in the exper-
iments conducted by Hermann & Rubatto (2009). In complex systems, epidote
group minerals incorporate REE primary through the following exchange reac-
tion: LREE+Fe2+ → Ca+Al (Giere & Sorensen, 2004; Armbruster et al., 2006).
However, in the studied system the absence of Fe may limit LREE incorporation
and probably inhibits crystallisation of real allanite.
1.4.3 Comparison with previous experimental studies of
monazite solubility
1.4.4 The general form of monazite solubility equation
Crystallisation/dissolution of monazite in the melt occurs by the reaction (Rapp
& Watson, 1986):
2(REE)PO4mnz ⇐⇒ REE2O3melt+ P2O5melt
(1.1)
The equilibrium constant of this reaction is expressed as:
Keq =[(REE)PO4]
2mnz
[REE2O3]melt[P2O5]melt
Where brackets [ ] denote the activity of the components and Keq is for the
equilibrium reaction constant. According to thermodynamic basics:
ln Keq =ΔG
RT≈ ΔU −ΔST + PΔV
RT∝ a + b/T + cP/T + dM (1.2)
where R is the gas constant, and T the absolute temperature, P is pressure,
ΔG is the standard Gibbs energy change, ΔU is the internal energy change, ΔS
is the entropy change and ΔV is the volume change of the reaction; a, b and c
are empirical parameters which can be derived from the experimental data by
linear multicomponent regression. Melt composition probably affects monazite
solubility through the change of the activity of LREE in the melt, but details of
this dependence are unknown. Therefore, the effect of melt composition usually is
approximated assuming a linear contribution from the compositional parameter
M with the coefficient d in the equation above.
30
Phosphorus is the fundamental component of all phosphate minerals, whereas
LREE are essential structural components (ESC Hanson & Langmuir (1978))
of monazite. The activity of LREE in monazite can be expressed by the mole
fraction of LREE in monazite XLREEmnz . Therefore the equation can be transformed
to:
ln K ≈ ln[(REE)PO4]mnz
[REE2O3]melt
≈ lnXLREE
mnz
LREEmelt
= ln XLREEmnz − ln LREEmelt ∝ a + b/T + cP/T + dM (1.3)
For convenience, all the parameters can be moved to the right hand side of the
formula:
ln LREEmelt = a + b/T + cP/T + dM + ln XLREEmnz (1.4)
This general formula of the LREE solubility in melt coexisting with mon-
azite contains a term for temperature, pressure, as well as melt and monazite
composition.
Solubility in granitic melts
The data from this study can be compared with earlier experimental studies
of monazite solubility (Rapp & Watson, 1986; Rapp et al., 1987; Montel, 1986,
1993). Data in this chapter can be directly compared with that of Rapp & Wat-
son (1986), as their natural monazite had a composition similar to the monazite
used in my experiments, and the melt composition is similar as well. In compar-
ison to the experiments of Rapp & Watson (1986), the results at 10 kbar show
very similar monazite solubility at 1200oC (Fig. 1.4, 1.9). At 1000oC and 10
kbar, I obtained a solubility of∑
LREE ≈ 3000 ppm, while Rapp & Watson
(1986) obtained 1880–1470 ppm. Somewhat higher solubility in my experiments
is probably attributed to the higher water and calcium content of the melt in my
experiments. In the 800oC experiments, the sum of LREE (La–Sm) is 344–624
ppm and experiments by Montel (1986, 1993) at similar conditions show 220–460
ppm Ce or Sm in peraluminous melts and 300-860 ppm Ce or Sm in metaluminous
melts.
31
Fig
ure
1.9:
New
calibra
tion
ofm
onaz
ite
solu
bility
inm
elts
and
its
com
par
ison
wit
hex
per
imen
taldat
aan
dpre
vio
us
calibra
tion
s.(a
)A
rrhen
ius
plo
tfo
rto
talLR
EE
solu
bility
inm
elts
.Sou
rces
ofdat
a:T
S–
this
study,
RW
–R
app
&W
atso
n(1
986)
,SB
–(Sko
ra&
Blu
ndy,20
10),
HR
–(H
erm
ann
&R
ubat
to,20
09)
and
R87
–R
app
etal
.(1
987)
(thes
edat
aar
enot
use
dfo
rca
libra
tion
ofeq
uat
ion
from
this
study).
Lin
esof
mon
azit
eso
lubility
calc
ula
ted
by
equat
ion
from
this
study
at10
wt.
%,fr
omeq
uat
ions
by
Rap
pet
al.
(198
7),M
onte
l(1
993)
(10
wt.
%H
2O
,D
=1)
and
Kel
sey
etal
.(2
008)
(FM
=1.
5).
Als
osh
own
mon
azit
eso
lubility
calc
ula
ted
from
form
ula
1.10
from
this
study
fordiff
eren
tpre
ssure
s.(b
)–
diff
eren
cebet
wee
nca
lcula
ted
and
exper
imen
talte
mper
ature
sfo
req
uat
ion
from
this
study.
Oth
erpar
amet
erar
eth
esa
me
ason
the
pre
vio
us
figu
re.
(c)
–th
esa
me
for
form
ula
by
Rap
pet
al.(1
987)
,(d
)–
the
sam
efo
rfo
rmula
by
Mon
tel(1
993)
.
32
In previous studies, several expressions for monazite solubility in granitic
melts were proposed and the general form for the solubility equation is (see
appendix):
ln LREEmelt = a + b/T + cP/T + dM + ln XLREEmnz (1.5)
Rapp & Watson (1986) performed dissolution experiments in a peraluminous
granite melt with low Ca, Fe and Mg (CaO+FeO+MgO = 1.3 wt.%) and 1–6
wt.% H2O at 8 kbar, over a temperature range from 1000 to 1400oC. Rapp
et al. (1987) presented an equation for the temperature dependence of monazite
solubility from experiments by Rapp & Watson (1986), which can be expressed
as follows:
ln∑
LREE = 18.5− 14160
T(1.6)
where∑
LREE is the sum of LREE (La to Sm) dissolved in the melt (in
ppm), and T is temperature in degrees Kelvin. The original equation by Rapp
et al. (1987) used weight fraction for LREE solubility, whereas in this study I
use ppm weight, resulting in a different value for the free term (18.5) in the
equation. This equations does not take into account any variations in solubility
as a function of melt and monazite composition.
Effect of Th
New experiments show significantly higher monazite solubility in high pressure
melts than those reported by Hermann & Rubatto (2009). An important dif-
ference between my experiments and those of Hermann & Rubatto (2009) and
Skora & Blundy (2010) is the much lower content of Th and U relative to LREE
in the starting material used in this study. Th and U are compatible in monazite,
and effectively act as additional REE, thereby suppressing the solubility of LREE
in the melt. In order to take into account the effect of Th and U, and to allow
for comparison of different experiments, LREE contents in the melt should be
corrected for the LREE activity in the melt. Montel (1993) suggested that the
activity of LREE in the melt could be constrained by the mole fraction of LREE
in monazite:
XLREEmnz =
LREEmnz
LREEmnz + Ymnz + Thmnz + Umnz
(1.7)
33
If the influence of Th is taken into account, then the different studies show
much better agreement with each other. Monazite solubility in the experiments
of Skora & Blundy (2010) is 172–573 ppm LREE at 800oC, 30 kbar; after con-
verting to LREE activity in monazite (XMnzREE=0.65–0.8), this rises to 215–820
ppm LREE, comparable to the results of this study. In the experiments of Her-
mann & Rubatto (2009) monazite composition was not determined. Calculat-
ing the composition of monazite in their experiments using partition coefficients
from this study shows that their monazite should contain XMnzTh =0.45–0.54 with
XMnzLREE=0.42–0.5. The composition of monazite calculated using the partition
coefficients from this study has less U than Skora & Blundy (2010) calculated
by mass balance, probably due to the presence of an additional U phase in the
experiments of Hermann & Rubatto (2009). At 800oC, 35 kbar, the solubility of
monazite measured by Hermann & Rubatto (2009) is 150 ppm, but their corrected
concentration is ≈300 ppm, comparable to experiments in this study. Therefore
if the influence of Th concentration in high-pressure, doped experiments is con-
sidered, then reasonable agreement between different studies is obtained.
Effect of pressure
This is the first study that systematically investigates the pressure dependence
of monazite solubility. In high pressure experiments in this study, monazite
solubility decreased with increasing pressure. In a number of these experiments,
the melt is enriched in both water and alkali elements relative to the melt in
reference experiments at 10 kbar, due to the crystallisation of silicates at high
pressure. These factors can increase monazite solubility (Montel, 1986; Rapp
et al., 1987) and mask the effect of pressure. However several of the experiments
run at different pressures have exactly the same melt composition (Fig. 1.4) and
all high pressure experiments show decreased monazite solubility in comparison
with the 10 kbar experiments.
Effect of melt composition and water content
In experiments by Rapp et al. (1987) performed at 8 kbar and 850–1000oC,
melts had a significant amount of CaO, FeO and MgO (CaO+FeO+MgO up to 7
wt.%). Monazite solubility appeared to be independent of composition in melts
with M = (2Ca + Na + K)/Si ·Al < 1.3, but was shown to increase rapidly
in melts with M > 1.3 (Fig. 1.10). The compositional dependence of monazite
saturation was expressed by the following relationship:
34
Figure 1.10: Effect of water content and melt composition on monazite solubilityin granitic melts. Data labels are the same as on figure 1.9, additionally M is forexperiments from Montel (1986, 1993) and R is for experiments by Rapp et al.(1987). * — denotes experiment which were not included into the calibrationof monazite solubility in this study. (a) The effect of water content can besatisfactorily described by square root dependence. Lines show dependence ofmonazite solubility from water content calculated by equations from this study fordifferent temperatures and by Montel (1993). (b) Monazite solubility increaseswith strong alkalinity Montel (1993) and with high content of Ca, Mg and FeRapp et al. (1987), however over the range of melt compositions obtained in thisstudy the effect of melt alkalinity on monazite solubility is insignificant. ThereforeI omit the compositional effect in the equation, but it is applicable only to meltswith low content of CaO, MgO, FeO and ASI > 0.85
35
ln∑
LREE = 8.27 + 0.84 ·S − 21.3 ·Si (1.8)
where S is the molar ratio (Na+K+Li+2Ca)/Al.
Montel (1986, 1993) performed experiments at 2kbar, 800oC, in the system Si-
Al-K-Na, using granitic compositions that ranged from peraluminous to peralka-
line under water saturated-conditions. These experiments showed that monazite
solubility increased dramatically in alkaline melts (Fig. 1.10). Montel (1993)
presented the following equation for calculating monazite saturation that takes
into account the effects of water content and melt composition:
lnREEt
XREEPO4
= 9.50 + 2.34D + 0.3879√
H2O − 13318
T(1.9)
where
REEt =∑ REEi(ppm)
at.weight(gmol−1)
and
D =(Na + K + Li + 2Ca)
Al· 1
(Al + Si)
Na, K, Li, Ca, Al, Si are in atomic fractions (with the total normalized to
unity); H20 is in wt.%; and XREEPO4 is the sum of the mole fraction of REE-
phosphates in monazite. The equation of Montel (1993) was calibrated using data
from three experimental studies (Rapp & Watson, 1986; Montel, 1986, 1993), but
it does not agree well with the dataset from Rapp et al. (1987), a discrepancy that
was attributed by Montel (1993) to the presence of Ca, Fe and Mg components
in the melts in the experiments of Rapp et al. (1987).
New data from this study allow also a better assessment of the influence of
high water contents on the LREE solubility in melts. The experiments of Rapp &
Watson (1986) were confined to water contents of 1–6 wt.%, whereas in this study
they go up to 17 wt.% H20. The compilation of LREE solubility as a function
of water contents (Fig. 1.10a) shows clearly that at high water contents, the
solubility increases less than at low water contents. New data are in agreement
with the suggestion of Montel (1993) that the effect of water content on the LREE
solubility in melts can be satisfactorily described by the square root dependence.
1.4.5 New expression for monazite solubility
In order to develop an equation that can predict monazite solubility as a function
of pressure, temperature as well as monazite and melt composition, I combined
36
new data with published studies and looked for the best linear equation to de-
scribe them. In addition to 19 new experiments from this study, there are 64
experimental runs published in which melts are buffered by monazite. LREE
concentrations were corrected to LREE activities in monazite, which take into
account the role of Th. Attempts to describe all experiments by a single equa-
tion were not successful: even if two compositional members were added to the
formula, the fit of the regression to the data was still poor (R2 = 0.81), and
the disagreement between the predicted solubility and the measured solubility
was unacceptably high. Comparison with experimental studies which considered
the effect of melt composition on monazite solubility (Montel, 1986, 1993; Rapp
et al., 1987) shows (Fig. 1.10) that over the range of compositions considered
in this work, the effect of melt composition on solubility is insignificant. How-
ever in more alkaline melts and/or in melts with high Ca, Fe and Mg content,
monazite solubility can increase substantially. Therefore I decided to limit the
calibration dataset to the experiments by Rapp & Watson (1986); Hermann &
Rubatto (2009); Skora & Blundy (2010) and experiments with peraluminous to
metaluminous granitic melts (ASI in the range 0.85–1.2) by Montel (1986, 1993).
57 experimental runs were used for the new calibration of monazite solubility in
melts. These experiments cover the range 2–50 kbar, 750–1400 oC, 1–20 % H2O
and 0.82–1.36 ASI. The effect of water content on the solubility can be approx-
imated by a square root dependence (Montel, 1993) (Fig. 1.10). The selected
experiments can be described by the following well-fitted (R2 = 0.95) equation:
ln∑
LREE = 16.16(±0.3) + 0.23(±0.07)√
H2O
−11494(±410)/T − 19.4(±4)P/T + ln XLREEmnz (1.10)
Where H2O is in weight percent, T is in Kelvin, P in kbar and∑
LREE
is the sum of La–Sm in ppm; XLREEmnz has been defined previously (Eq. 1.7).
One standard deviation uncertainties on parameters are given in parentheses.
The equation reproduces the experimental data with absolute average deviation
of 0.19. Addition of various compositional parameters (ASI = Al2Ca+Na+K
, D =Na+K+Li+2Ca
Al1
Al+Si, M = Na+K+Li+2Ca
AlSi, CNKM = Ca
Na+K+Mg+Ca, etc.) to this for-
mula does not produce any improvement in the fit, since their coefficients were
statistically insignificant. Therefore I conclude that for the selected set of ex-
periments, the effect of melt composition on monazite solubility is insignificant.
On the other hand, this equation should probably be used exclusively for pera-
37
luminous and metaluminous compositions. This limitation is tolerable, because
peraluminous melts are much more common than peralkaline melts. The pro-
posed equation is suitable for peraluminous and metalumious (with ASI >0.85)
leucogranitic melts (with CaO+FeO+MgO<3 wt.%). Caution is needed if the
equation is used outside of the range of conditions and melt compositions on
which it was calibrated.
The graphic representation of the published calibrations of monazite solubil-
ity, including the newly proposed one, are shown on Fig. 1.9. Calibrations by
Rapp et al. (1987) and Montel (1993) are in perfect agreement with each other
at D=1, which is not surprising, given that their temperature dependencies are
based on the same experiments. The calibration for monazite solubility predicts
monazite saturation at 100 ppm LREE in melt at 700oC, 10 kbar and 5 wt.%
H2O. In compositions used for calibration of the equation 1.10 parameter D
varies from 0.9 to 1.4 and and for this range of melt compositions (and at 700oC and 5 wt.% H2O) the equation by Montel (1993) predicts monazite satura-
tion at 40–130 ppm LREE, consistent with new calibration. This calculation
demonstrates large effect of melt composition on the monazite solubility in cali-
bration by Montel (1993), which is not observed on the larger dataset (Fig. 1.10).
Therefore new calibration is more suitable for predicting monazite saturation in
peraluminous and metaluminous melts , whereas the equation of Montel (1993)
is better suited to peralkaline–metaluminous melts.
1.4.6 Monazite/melt partitioning
Monazite/melt partition coefficients have been estimated before using natural
samples (Yurimoto et al., 1990; Ward et al., 1992; Bea et al., 1994), based upon
bulk rock and monazite compositions. Data for Kmnz/lD REE/Ce(1000) at 10 kbar,
1000oC are in excellent agreement with values from natural samples (Yurimoto
et al., 1990; Ward et al., 1992) (Fig. 1.5b). There is a significant difference be-
tween data from this study and monazite/melt partition coefficients calculated
by Bea et al. (1994), which have a much higher D for the HREE than that deter-
mined in this study. This disagreement might be due to the fact that Bea et al.
(1994) measured monazite/bulk leucosome partitioning; because leucosomes can
contain peritectic phases from the melting reactions, their composition can differ
significantly from the parental melt.
Monazite/melt partition coefficients for La, Ce, Nd, Sm, Th and U were
determined experimentally at 800–900oC and 30 kbar by Skora & Blundy (2010).
38
They observed that Dmnz/lLa -D
mnz/lNd are equal within analytical precision, that
Dmnz/lLREE is 20–30% higher than D
mnz/lSm , and that D
mnz/lTh is 16–24 times greater
than Dmnz/lU . All these observations are consistent with the dataset. However,
a wide variations in Th/La ratios (Kmnz/lD Th/La), were observed, in contrast to the
constant ratios seen in experiments in this study.
Dmnz/lLa−Nd obtained in this study are constant. This effect is not due to the
specific ratio of LREE in the starting material, because in the experiments the
relative proportions of LREE are close to chondritic (or natural). In the experi-
ments by Skora & Blundy (2010) all the REE had almost identical concentrations,
but similar Dmnz/l for LREE (from La to Nd) were obtained. This behaviour of
LREE in monazite is somewhat different than in other REE minerals, which
usually show a preference for a particular rare earth. For instance zoisite/melt
partitioning of REE has a peak at Nd (Frei et al., 2003), and allanite/melt par-
titioning peaks at La (Klimm et al., 2008). The observed partition coefficients
for LREE can not be satisfactorily fit to the parabolic dependence with respect
to ionic radius (Onuma plot) for the LREE (Blundy & Wood, 1994) (Fig. 1.11).
However, the HREE do fit the parabolic form. Therefore, monazite provides
a unique example of deviation from the Onuma model. The inconsistency of
monazite/melt partitioning with the model’s predictions is attributable to the
fact that the LREE constitute essential structural constituents (ESC Hanson &
Langmuir (1978)) in monazite, and therefore would not be expected to behave
according to Henry’s Law. In addition, some phosphates (apatite in particular)
are known to have a flexible structure that can accommodate cations of different
size (White & Dong, 2003).
This study shows that Kmnz/lD HREE/Ce(1000) increases with temperature and de-
creases with the water content of the melt (Fig. 1.6). It is unlikely that these
variations will be important in controlling HREE behaviour in natural systems,
because these elements are controlled not by monazite but by other minerals
(garnet, first of all). The data show a weak correlation of Kmnz/lD Sm/Ce(1000) with
temperature and probably pressure. Partition coefficients for the REE begin to
decrease with Sm, and because the partitioning of HREE is variable, it is entirely
possible that the extent of fractionation of Nd from Sm also varies with temper-
ature and/or pressure. The experiments show much lower Dmnz/lU than D
mnz/lTh
and an increase of U compatibility in monazite with pressure and temperature.
Because the experiments were performed at quite reduced conditions (evident
from negative Eu anomaly) it is likely that U in melts was present as U4+. Un-
der more oxidized conditions some U will be present as U6+ and thus Kmnz/lD Th/U
39
Figure 1.11: Onuma diagram for monazite/melt partitioning. Black dots repre-sent 95% confidence intervals.
in monazite will is expected to be even higher than in experiments in this study.
The Th/La ratio of magmatic rocks and sediments is an important geochem-
ical indicator (Plank, 2005). In natural monazite, ThO2 content is 4–12 wt.%
(Forster, 1998); in monazite from amphibolite facies rocks and in granites the
cheralite substitution (Ca0.5Th0.5PO4) often predominates (Pyle et al., 2001;
Forster, 1998). In experimental monazite, the extent of cheralite and huttonite
substitutions are approximately equal. The incorporation of Th into monazite
can occur by either one of the following reactions:
CePO4 + ThO2 + SiO2 ⇒ (Ce, Th)(PO4, SiO2)
2CePO4 + ThO2 + CaO ⇒ ThCa(PO4)2 + Ce2O3
These reactions show that the ThO2 content of monazite can be affected by the
activities of Ca, Si and P, which depend on melt composition and are expected
to vary significantly from case to case. The Th/La of experimental melts in
this study display only minimal fractionation of the Th/La ratio from monazite
and/or the starting composition (Fig. 1.7), and the CaO and SiO2 contents of
the anhydrous melts vary only from 1.9 to 3 wt.% and from 72.2 to 76.6 wt.%,
respectively. The experiments cover a narrow range of compositions and it is
40
still possible that there is a dependence of Kmnz/lD Th/La from melt composition, but
within the range of liquids studied, data do not allow this to be ascertained with
any confidence.
Skora & Blundy (2010) observed a large fractionation of Th from La in ex-
perimental melts coexisting with residual monazite at 30 kbar, 700–900oC, in
contrast to the findings of this study. They argued that this could be related
to high Th contents in doped experiments. In such a case, Th is a major con-
stituent and Th-LREE solid solutions are governed by Raoult’s Law, whereas
in natural compositions Th is a minor component and Th-LREE solutions are
governed by Henry’s Law. Consequently, Skora & Blundy (2010) concluded that
D(Th) will depend on the Th/LREE ratio of the system and may be subject to
large variations. The suggestion that Th can affect monazite partitioning sys-
tematics is reasonable. For example, the amount of HREE soluble in monazite
in equilibrium with xenotime increases as the Th content in monazite increases
(Seydoux-Guillaume et al., 2002). A Th-doped starting material was also used in
the experiments of Hermann & Rubatto (2009), where high Th monazite coex-
isted with hydrous melts. Although the present set of experiments used a doped
starting material, I kept the Th/LREE ratio similar to upper crustal rocks. As a
result, X(Th) in my experiments is 0.07 — close to the amount typically found in
natural monazite (Forster, 1998). My experiments indicate that at such relatively
low X(Th) the fractionation of Th from La between monazite and melt is minimal
over the whole range of experiments (Fig. 1.6, 1.7). Skora & Blundy (2010) also
proposed that large variations in the Th/La ratios of their experiments might be
due to the effect of H2O content on the partitioning of LREE and Th. However,
new data show that water content has little effect on Dmnz/meltTh .
Monazite and allanite are the two primary LREE-bearing minerals in both
felsic and metabasic rocks (Bea, 1996; Klimm et al., 2008). Trace element parti-
tioning behaviour in the presence of monazite and allanite is in general similar,
with both minerals showing a preference for LREE relative to HREE and Th
relative to U (Klimm et al., 2008), although there are some differences: allanite
shows a preference for La, whereas monazite partitions all the LREE equally.
Also, Th is more compatible in monazite than the LREE (e.g., La–Ce), whereas
it is less compatible than the LREE in allanite.
Arsenic most likely substitutes for P in the monazite structure, as there is
the arsenic equivalent of monazite — gasparite LREEAsPO4 (Ondrejka et al.,
2007).
Natural monazites typically show zoning with respect to HREE, Y, Th and U
41
as well as SiO2 and CaO. In this study, I found that temperature, pressure and
melt composition can all have an effect on monazite-melt partitioning. However, I
should emphasize that variations in partitioning occurred over an extremely large
pressure-temperature range, and yet monazite/melt partitioning was remarkably
constant at conditions relevant to the continental crust. It is unlikely, therefore,
that variations in monazite-melt partitioning can be responsible for the compo-
sitional variation observed in natural monazite from high grade rocks. These
variations rather must originate from changes in the overall phase assemblage
and/or the bulk composition of the system.
1.5 Implications of the experimental results
1.5.1 Implications for fractional crystallisation of granites
Because monazite is the major host for LREE in granites (Bea, 1996), its frac-
tionation can have a significant effect on the REE pattern of its granite host
(Wark & Miller, 1993). Using experimental partition coefficients from this study,
I modeled Rayleigh fractionation of monazite from a typical granitic melt com-
position (Fig. 1.12). The calculations predict enrichment of the melt in HREE
relative to LREE, and in U relative to Th. LREE from La to Nd are not fraction-
ated by monazite crystallisation from granitic melts. Therefore the formation of
monazites enriched in specific LREE (Nd-monazite and Sm-monazite) cannot be
related to the fractionation of magma but rather indicates the effect of a fluid
phase. If a significant degree of monazite fractionation has occurred, then the
REE pattern can appear to be depleted in Nd relative to Pr and Sm, e.g. a
neodymium anomaly develops. Also, monazite fractionation causes a significant
increase in the Sm/Nd ratio. Data from this study provide some evidence for a
decrease in Kmnz/lD HREE/LREE with decreasing temperature, so that in low-T melts,
monazite fractionation will have a stronger effect on the composition of the melt.
There are occasional reports of natural granites with chondrite-normalized
REE patterns showing notable depletions in specific elements (e.g., Nd-depleted
granites, fractionated Y/Ho and Zr/Hf ratios, etc.); these REE patterns have
been attributed to the “tetrad effect”(Bau, 1996; Yasnygina & Rasskazov, 2008),
or to monazite fractionation (Yurimoto et al., 1990). The concept of the “tetrad
effect” was adopted from marine geochemistry (Bau, 1996); the REE are sub-
divided into 4 groups of 4 elements each (hence the term tertrad), with only
slight variations within groups but abrupt changes between them. It was pro-
42
Figure 1.12: Modeled trace element composition a granitic melt with fractionalcrystallisation of monazite. Numbers show the percentage of monazite crys-tallised.
posed that the effect arises from ions complexing with volatiles (H2O, Li, B, F,
P, Cl) (Bau, 1996), and in particular from extraction of fluorine rich fluid from
melt (Veksler et al., 2005). New experimental data and calculations (Fig. 1.12)
confirm the conclusion of Yurimoto et al. (1990) that a Nd anomaly can be pro-
duced by monazite fractionation. However, monazite fractionation is unlikely to
change the La/Ce ratio, or to decouple Y from Ho, which are both common fea-
tures of evolved granites showing the ”tetrad effect” (Irber, 1999; Veksler et al.,
2005). Therefore, monazite fractionation is probably unable to account for all the
features of ”tetrad effect” granites and other explanations, such as partitioning
in the presence of high F melts/fluids (Veksler et al., 2005), are better able to
account for the origin of these unusual melts.
Cuney & Friedrich (1987) observed that the nature of the mineral hosting
U in granites commonly depends on whether monazite or allanite are the pri-
mary LREE-bearing phase: U is concentrated in uraninite (UO2) in granites
with monazite, but in high-Ca granites with allanite, uranothorite ((U, Th)O2)
is commonly present. This can be explained by differences in the partitioning
behaviour of monazite and allanite: residual melts after allanite or monazite crys-
tallisation will have elevated U concentrations, but residual melts after monazite
crystallisation will also have low Th contents, favoring precipitation of urani-
nite. Allanite fractionation causes enrichment in both Th and U and subsequent
crystallisation of uranothorite. Therefore monazite crystallisation will have a
43
profound impact on the trace element signature of the granitic host melt, and on
its accessory mineral assemblage.
1.5.2 Implications for crustal melting
Monazite is one of the main hosts for LREE, Th, and U in migmatites, and is a
common residual phase during crustal anatexis (Bea, 1996). Monazite solubility
in melt is a critical parameter for understanding the behaviour of these elements
during partial melting in the crust. If the amount of monazite dissolved in the
melt is lower than the bulk rock content of LREE, then melt extraction will
cause LREE enrichment in the restite (Rapp et al., 1987). On the other hand,
if monazite solubility in the melt during anatexis is significantly higher than its
content in the bulk rock, then melt extraction will lead to depletion of the residue
in LREE, Th and U. In the extreme case where all accessible monazite is dissolved
in the melt, severe depletion of the residue in LREE, Th and U will occur. In this
case the main host for these elements will be the melt, and their concentrations
in the residue will be controlled by the partitioning behaviour of the remaining
residual minerals with respect to the melt.
The behaviour of monazite and zircon during partial melting of metapelites
was modeled by Kelsey et al. (2008), using phase relationships for silicates de-
termined by THERMOCALC. Zircon and monazite solubilities were calculated
using original expressions derived from experimental data. The modeling predicts
that monazite should be completely dissolved in the melt in the early stages of
anatexis, and at much lower temperatures than zircon. However this prediction
is inconsistent with the commonly observed preservation of LREE in restites fol-
lowing anatexis (e. g. Schnetger (1994)). The expression for monazite solubility
in the model of Kelsey et al. (2008) was derived by regression of the experimental
data of Rapp et al. (1987) expressed as:
lnLREEmnz
LREEmelt
= −310
T− 1.324 ·FM + 7.5852 (1.11)
Where FM is the molar ratio (K+Na-2(Ca+Mg+Fe))/(Al*Si) and T is tem-
perature in Kelvin. One particular feature of this expression is the very weak tem-
perature dependence (in fact this equation predicts even a decrease of monazite
solubility with temperature increase) and high monazite solubility at 700-800oC.
The experiments of Rapp et al. (1987) were designed to ascertain the role that
melt composition play in monazite solubility. In fact the authors purposely state
that “the effect of temperature is not discernible in the monazite solubility data”,
44
emphasizing the fact that the temperature dependence of monazite solubility in
the range 850–1000oC was in the same range as the resolution of the analyti-
cal technique (EPMA), being employed at the time. Equation 1.11 from Kelsey
et al. (2008) predicts ≈1950 ppm of LREE in melt at T=700oC and FM=1.2
(Fig. 1.9). This very high solubility led Kelsey et al. (2008) to conclude that
monazite would be fully dissolved from a source rock with a bulk LREE content
of 170 ppm at temperatures as low as ≈ 700oC. This is unrealistic, because many
migmatites preserve most of their REE after extraction of partial melts (e. g.
Rollinson & Windley, 1980; Schnetger, 1994). Because partial melting of crustal
rocks produces broadly granitic melts, the principle variable responsible for the
LREE solubility is the temperature of melting and not the composition of the
melt (Figs. 1.9, 1.10). Therefore the calibrations of Rapp et al. (1987), Montel
(1993) and this study, which estimate monazite solubility in peraluminous melts
at 700oC at 50–100 ppm, would be more suitable for modeling the behaviour of
monazite during anatexis. I propose that because of the choice of experiments
for the calibration of monazite solubility, the prediction of early dissolution of
monazite during melting in the model by Kelsey et al. (2008) is inconsistent
with observational data. Therefore the conclusion of Kelsey et al. (2008) that
in migmatites metamorphic monazite has younger ages than zircon due to the
difference in solubility probably needs to be reassessed.
The fractionation of LREE, Th and U is ultimately controlled by the overall
partitioning of elements amongst the residual phases and anatectic melt, as well
as the proportions of phases and the degree of melting. In addition to monazite,
other accessory phases such as allanite, zircon and apatite, and major minerals
such as garnet, play an important role in storing these elements. Therefore,
in order to calculate the trace element signature of a partial melt buffered by
monazite one would need to consider all these phases. Such a task is beyond the
scope of this paper; nevertheless, some preliminary conclusions can be made from
qualitative application of new partitioning values.
The LREE have the same monazite/melt partition coefficients. Therefore, if
partial melting is buffered by monazite, REE ratios like La/Nd should not change
unless another mineral contains a significant fraction of LREE. The behaviour of
the MREE and HREE is more complex, because there are many phases that can
host these elements.
Thorium is slightly more compatible in monazite than the LREE, with Kmnz/lDCe/Th
approximately 1.3. Very low degree anatectic melts in equilibrium with monazite
will have a 30% lower Th/La ratio than the source rock. As the degree of melting
45
increases, the Th/La ratio of the melt will increase approximately linearly (be-
cause Th is slightly more compatible than LREE); when monazite is completely
dissolved, the Th/La ratio of the melt will be that of the original source. In
the residue, the situation will be the exact opposite. At low degrees of melting,
Th/La ratios will be relatively unaffected, and the last residual monazite will
have a Th/La ratio that is 30% higher than the melt and the source. The great-
est degree of fractionation of Th relative to the LREE that can be produced by
one stage melting in the presence of monazite is limited to approximately 30%.
This effect is quite small, and can easily be masked in nature by variability in the
composition of the source rock’s precursors. On the local scale these fractionation
effect can be more pronounced. Monazites with Th-rich cores are often observed
in granulites after partial melting (Zhu & O’Nions, 1999). These monazite cores
may represent residual grains from peak melting conditions, with the low Th rims
formed by crystal overgrowth from unsegregated partial melt.
Many granulites, metapelites and metafelsic rocks are strongly depleted in U
relative to their low grade equivalents (Rudnick et al., 1985; Nozkhin & Turkina,
1993; Schnetger, 1994; Bea & Montero, 1999). This feature has been explained
by uranium extraction either by fluids or by partial melts. Th is almost insoluble
in fluids and U solubility could be very high in oxidized fluids with high chlorine
content (Keppler & Wyllie, 1990; Bali et al., 2011). A large fraction of U in
many metamorphic rocks is stored in monazite, but monazite/melt partition
coefficients for U are much lower than for Th and the LREE. Therefore anatectic
melts buffered by monazite should be enriched in U relative to Th and LREE, and
their extraction will deplete the residue in U relative to Th and LREE. This is in
agreement with the presence of peritectic melt inclusions with high U contents
and low Th/U ratios (Acosta-Vigil et al., 2010).
1.5.3 Implications for melting in subduction zones
Experimental studies have highlighted that fluid fluxed partial melting of sed-
iments is a key process in the mobilization of incompatible elements from the
subducted slab (Hermann & Green, 2001; Hermann & Spandler, 2008; Skora &
Blundy, 2010). It has recently been proposed that monazite may be an ubiquitous
residual phase during melting of metasediments in subduction zones, attracting
considerable attention due to monazite’s ability to affect key element ratios such
as Th/La and Th/U (Plank, 2005; Hermann & Rubatto, 2009; Skora & Blundy,
2010).
46
The similarity of Th/La in subducting sediments and in the subduction com-
ponent of arc magmas led to the proposition that Th/La is not significantly frac-
tionated during melting of metasediments (Plank, 2005). Because allanite can
fractionate LREE from Th, it was proposed that the Th/La ratio is controlled by
residual monazite (Plank, 2005). This has been confirmed by the experiments of
Hermann & Rubatto (2009), albeit for compositions that contained an amount
of Th significantly higher than natural sediments. On the other hand, Skora &
Blundy (2010) observed large variations in Th/La in their experimental melts in
equilibrium with residual high-Th monazite. In order to produce a melt with a
similar Th/La as the input sediments, as inferred for sub-arc melting, Skora &
Blundy (2010) proposed that monazite is unlikely to be a residual phase in sub-
ducted sediments that have undergone partial melting. Instead they suggested
that the continuous flushing of externally derived water trough the sediments
would cause melting sufficient to completely dissolve all monazite present in the
source rock. New data now show that the fractionation of Th from La imposed
by monazite is only ≈30%.
Therefore, the presence of monazite in the residue can produce a maximum
30% decrease in the Th/La ratio of the subduction component of arc magmas.
Allanite will have the opposite effect on the Th/La ratio because Th is more in-
compatible than La in this phase (Hermann, 2002a; Klimm et al., 2008; Hermann
& Rubatto, 2009). The Th content of monazite can potentially affect the Th/La
ratio of coexisting melts (Skora & Blundy, 2010). The Th/La ratio of GLOSS
is ≈0.24 (Plank & Langmuir, 1998), and the molar ratio Th/(LREE+Th) is
≈0.05. Considering that over 95% of the LREE and Th under subsolidus con-
ditions is hosted by either accessory allanite or monazite (Hermann, 2002a), the
molar Th/(LREE+Th) in the accessory phases is the same as in the bulk rock.
Here experiments were conducted at a slightly higher Th/(LREE+Th) of ≈0.07,
representative of Th-rich subducted sediment compositions (Plank & Langmuir,
1998). Therefore, new results suggest that residual monazite in the subducted
sediments will not significantly (<30%) affect the Th/La of the melt, and there is
no need for monazite to be completely consumed in order to transfer the Th/La
ratio of sediments into arc magmas. Melts produced in subducted sediments with
residual monazite are in agreement with the observed systematics presented by
Plank (2005).
The presence of residual monazite during partial melting of subducted sedi-
ments is a key assumption for the application of a “slab thermometer”. Plank
et al. (2009) proposed that the H2O/Ce ratio of primitive arc melts can be used
47
to estimate temperature at the top of the slab. This thermometer is based on
the presence of residual monazite during metasediment melting, and the strong
temperature-dependence of monazite solubility. Aqueous fluids will have ex-
tremely high H2O/Ce as water contents are very high and Ce contents are very
low. Low temperature melts from metasediments are water-rich and monazite
solubility will be low, and consequently these melts will maintain high H2O/Ce
ratios. A further increase in temperature will result in an increase in the LREE
content and decrease in the water content of the melt, and a lowering of H2O/Ce.
Primitive melt inclusions in olivine display H2O/Ce that indicate sediment melt-
ing at 700–900oC Plank et al. (2009). new data confirm a strong temperature
dependence monazite solubility at high pressure. At pressures of 30–50 kbar,
monazite solubility is 30–40% lower than at 10 kbar. Bulk Th contents also af-
fects monazite solubility, resulting in lower LREE contents for melts produced
in experiments with high Th contents such as those of n Hermann & Rubatto
(2009) and Skora & Blundy (2010). New expression for monazite solubility in
granitic melts (equation 1.10) includes a pressure term, a water term and a Th-
composition term, and allows for a more accurate assessment of the H2O/Ce
ratio in granitic melts in subduction zone in the presence of residual monazite.
A typical feature of arc melts is their low Th/U ratio (Taylor & McLennan,
1985; Hawkesworth et al., 1991; Plank, 2005). In arc magmas, the Th/U ratio
is 1.5–3.0, whereas the Th/U ratio of the primitive mantle, the bulk Earth, the
upper continental crust and most sediments is around 4 (Taylor & McLennan,
1985). The enrichment of arc magmas in U is proposed to be either due to
fluids derived from metabasites (Kelley et al., 2005; Becker et al., 2000) or from
melting of metasediments (Klimm et al., 2008; Hermann & Rubatto, 2009). In the
latter case, preferential partitioning of Th relative to U in allanite and monazite
could account for melts enriched in uranium. I observed that with increasing
pressure, the partitioning of U into monazite increases significantly, although
Th remains 5 times more compatible than U. Therefore, despite the increased
compatibility of U in monazite with increasing pressure, melts in equilibrium
with monazite will retain low Th/U, supporting previous models (Hermann &
Rubatto, 2009) attributing the low Th/U of arc magmas to residual monazite in
subducted sediments.
48
1.6 Conclusions
Monazite has partition coefficients higher than unity for REE, Th, U, Y, As
and V, with high preference for LREE(La-Nd). Starting with Sm, partition
coefficients decrease, and for Lu they are 9–20 times lower than for the LREE.
Th is ≈30% more compatible in monazite than the LREE. Partitioning of U into
monazite increases with pressure, but under all experimental conditions studied,
U is less compatible in monazite than LREE and Th. Monazite/melt partitioning
for LREE is constant and does not fit the Onuma model.
Based on the new experimental data and previous studies, I derive the fol-
lowing equation for monazite solubility in leucogranitic melts:
ln∑
LREE = 16.16(±0.3) + 0.23(±0.07)√
H2O
−11494(±410)/T − 19.4(±4)P/T + ln XLREEmnz
Where H2O is in weight percent, T is in Kelvin, P in kbar and∑
LREE is
the sum of La-Sm in ppm; XLREEmnz is the molar ratio of LREE to the sum of all
cations (REE, Th and U) in monazite.
Monazite fractionation from granitic melt can cause a significant increase in
the ratio of LREE/HREE and Th/U, and in extreme cases, may produce negative
Nd anomalies in the REE patterns of fractionated granitic melts. The presence
of monazite as a residual phase will cause little fractionation of LREE relative to
Th, and significant depletion in U relative to these elements.
49
Table 1.1: Compositions of starting materials. *–with 5 wt.% of trace elements;**– water is added as Al(OH)3 and its content is calculated from stoichiometry.
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tco
mpos
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Chapter 2
Association of rocks with
different PT paths within the
Barchi Kol’ UHP block
2.1 Introduction
The Kokchetav metamorphic belt in northern Kazakhstan is famous because of
its ultra high pressure (UHP) metamorphic rocks, which host abundant meta-
morphic microdiamonds and other indicators of extremely high pressures. The
peak conditions are estimated to reach 45–70 kbar and 950–1000oC (Sobolev &
Shatsky, 1990; Chopin, 2003; Ogasawara et al., 2002) at 530–520 Ma (Claoue-
Long et al., 1991; Hermann et al., 2001). Numerous studies have shown that
the Kokchetav rocks were subducted to a depth of more than 120 km and then
exhumed to the surface (Dobretsov et al., 1995; Kaneko et al., 2000).
In the Kokchetav complex, metamorphism varies from low-P metamorphism
in the Daulet suite to extreme conditions in UHP rocks. Numerous works deter-
mined PT conditions of metamorphism in the various domains of the Kokchetav
complex (e.g. Sobolev & Shatsky, 1990; Theunissen et al., 2000; Chopin, 2003).
Studies of non-UHP rocks provided constraints on the general tectonic history of
this unique terrain (Buslov et al., 2010; Dobretsov et al., 2005, 2006; Zhimulev
et al., 2010).
Worldwide ultrahigh pressure metamorphic terrains are located within oro-
genic belts (Fig. 2.1). Tectonic studies of the Kokchetav complex concluded
that UHP rocks are surrounded by rocks with different peak PT conditions (Do-
bretsov et al., 1995; Kaneko et al., 2000; Liou et al., 2002). The UHP blocks
in these models are considered as coherent units with consistent peak PT condi-
tions across the whole unit and the same exhumation paths (Kaneko et al., 2000).
UHP rocks are usually observed in terrains which have tectonic contacts with ad-
59
60
Figure 2.1: Tectonic map of Eurasia (after Wang et al. (2011)), location of theKokchetav complex and other important UHP complexes.
jacent, lower pressure rocks (Kaneko et al., 2000; Dobretsov et al., 1995). These
terrains are composed of rocks with variable lithologies and different indicators
of UHP conditions, which poses several questions. Have all the rocks in the UHP
terrain experienced the same metamorphic path? In particular, have all rocks
achieved the same peak metamorphic conditions? Were these conditions homo-
geneous during exhumation? In this chapter I investigate the metasedimentary
rocks of the Barchi Kol’ UHP terrain in the Kokchetav massif. The petrology
and geochronology of these rocks were studied in detail in order to understand
the petrogenetic relations with other rocks of the Kokchetav complex. Miner-
alogical and geochemical features indicate that some of these rocks experienced
metamorphism at different peak conditions and had different PT paths, despite
their close spatial association with UHP rocks.
2.2 Geologic setting
2.2.1 Geology of the Kokchetav complex
The Kokchetav metamorphic belt (KMB, also known as Kokchetav complex or
massif) is located in northern Kazakhstan about 270 km northwest from Astana,
the new capital of Kazakhstan and 900 km west of Novosibirsk, a major city in
Siberia, Russia (Fig. 2.1). The geography comprises low hills and plains covered
by grassland steppe with many small lakes and small forests. Rock exposure is
61
relatively poor in the region.
The Kokchetav metamorphic belt is located inside the Central Asia fold belt,
which covers thousands of kilometers between the East European, Siberian, North
China and Tarim cratons (Fig. 2.1) (Zonenshain et al., 1990; Wang et al., 2011).
The Central Asia fold belt is comprises of Paleozoic continental crust, which is
itself composed of oceanic arcs, sedimentary complexes and blocks of precambrian
continental crust, which are separated by ophiolites, sometimes with high pressure
rocks (Buslov et al., 2009; Volkova et al., 2008) and intruded by abundant granitic
intrusions. The Kokchetav microcontinent represents one of these ancient blocks.
Arcs formed during the early Paleozoic oceanic (Caledonian orogeny, 490–390
Ma), were assembled during collisional events during the Hercynian–Variscan
orogeny (380–280 Ma) and development of the belt was completed in the Permian
(Zonenshain et al., 1990). The Kokchetav complex has an Early–Mid Cambrian
age (Claoue-Long et al., 1991; Hermann et al., 2001; Katayama et al., 2001) and
is surrounded by younger continental structures (Zonenshain et al., 1990).
The Kokchetav metamorphic complex has a belt shape that extends from east-
east-north to west-west-south over 150 km, with boundaries defined by younger
terrains. The southern boundary of the Kokchetav complex is comprised by the
Zerenda granite batholith. North of the Kokchetav complex occurs the North
Kokchetav tectonic zone, while further north the Stepnyak paleo-island arc crops
out. The North Kokchetav tectonic zone is composed of thrust sheets of low
grade rocks, lenses of HP rocks, and contains olistostrome formations which
demonstrate sedimentation during orogenesis (Zhimulev et al., 2010, 2011). The
Stepnyak paleo-island arc if composed of low grade sediments and voulcanic for-
mations. The North Kokchetav tectonic zone formed due to collision of the
Kokchetav microcontinent and the Stepnyak paleo-island arc in Early–Middle Or-
dovician, substantially later than Early–Middle Cambrian age of the Kokchetav
metamorphic complex.
The Kokchetav complex is composed of units with different metamorphic con-
ditions, which researchers have named suites, domains and/or terrains (Rozen,
1971). The main fraction of the complex is composed of low grade precambrean
sedimentary rocks, which are interpreted as sedimentary cover of the Kokchetav
microcontinent, and felsic gneisses exposed in several localities as its basement
(Turkina et al., 2011). Within these precambrean suits tectonically juxtaposed
HP and UHP terrains (from west to east): Barchi Kol’, Kumdy Kol’, Sulu Tjube,
Enbek Berlyk, and Kulet. UHP rocks are exposed in Kulet, Kumdy Kol’, Barchi
Kol’ and there have been reports of findings of diamond-bearing rocks further
62
West from the Barchi Kol’ area (Shatsky et al., 2005). Kumdy Kol’ and Bachi
Kol’ reached sufficiently high pressures to stabilize diamond and peak tempera-
ture are estimated at 950–1000oC, whereas in the Kulet area, metamorphism
occurred in the stability field of coesite at lower temperatures of 720–760oC
(Parkinson, 2000).
There are two general interpretations of the regional structure of the Kokchetav
metamorphic belt. One is the ”transpressional” or ”megamelange” model pro-
posed by Russian geologists (Dobrzhinetskaya et al., 1994; Dobretsov et al., 1995).
In this scenario, the Kokchetav complex represents a megamelange composed of
terrains of different metamorphic history bounded by normal faults (Dobrzhinet-
skaya et al., 1994; Dobretsov et al., 1995; Theunissen et al., 2000). It was proposed
that the western and eastern parts of the complex have different deformation pat-
terns and probably were exhumed by different mechanisms (Theunissen et al.,
2000).
Another interpretation is the ”subhorizontal model” proposed by Japanese
scientists (Kaneko et al., 2000). In this model, the primary structure of the
Kokchetav complex is subhorizontal and layered. The UHP rocks are enveloped
by layers of HP rocks. The HP-UHP sheets are underlied by low-P metamor-
phic rocks of the Daulet Suite (Buslov et al., 2010) and are overlain by weakly
metamorphosed to unmetamorphosed sedimentary strata (Kaneko et al., 2000).
These folded strata are divided by subhorizontal faults and bound by normal
faults. This model of subhorizontal structure of the KMB is based on good expo-
sures in the eastern part of the KMB, particularly the subhorizontal structures in
the localities of Sulu-Tjube and Solbat Kol’ (Kaneko et al., 2000; Ishikawa et al.,
2000). However, in the UHP areas such observations are absent, and usually the
UHP rocks dip at steep angles (Shatsky et al., 1995).
According Dobretsov et al. (1995) the Kokchetav UHP belt was then formed
due to the collision of the Kokchetav continent with the Stepnyak island arc.
This event formed one of the segments of the Kazakhstan continent, which in the
late Paleozoic was accreted and incorporated into the structure of the Central
Asia Fold belt. Also it was proposed that the Kokchetav UHP rocks were formed
in a Himalayan type collision and exhumation was triggered by slab-breakoff
(Maruyama & Parkinson, 2000).
63
2.2.2 Geology of Barchi Kol’ unit
The most famous UHP locality in the Kokchetav complex is Kumdy Kol’, where
exploration audit was excavated and UHP rocks are abundant. The Barchi Kol’
UHP unit is a terrain located 17 km west of the Kumdy Kol’ UHP unit near
Barchi Kol’ lake. It is elongated from southwest to northeast and has a size of
approximately 2.5 × 5 km (Fig. 2.2). The Barchi Kol’ UHPM unit is bound in
the northwest and north by faults separating the UHP rocks from the weakly
metamorphosed Precambrian sediments of Kokchetav and Sharyk suites. To the
south of this UHP unit, the Krasnomai alkali-utrabasic complex occurs, composed
of pyroxenites, micaceous pyroxenites and carbonatites (Letnikov et al., 2004).
Further south there is the ”White Lake domain” or ”Efimov suite” (Dobretsov
et al., 1995, 1998). Dobretsov et al. (1999) described it as ”amphibolites, amphi-
bole schists and quartzites, assumed to represent fragments of an oceanic crust”.
Further south there is the large Silurian-Devonian Zerenda Batolith composed of
granites and granodiorites.
The internal structure of the Barchi Kol’ unit is known from the mapping
and drilling carried out by the Kokchetav Prospecting Expedition (Fig. 2.2),
the Kokchetav geological survey and a surface map produced by Masago (2000).
The following rock types are described in the Barchi area: eclogites, garnet-
pyroxenites, amphibolites, calc-silicates, migmatites, schists and a variety of
gneisses. The most abundant rock type are gneisses, which contain feldspars,
quartz and garnet. Based on their subordinate mineral phases, gneisses can be
subdivided into kyanite, clinopyroxene, clinozoisite, biotite, and two-mica bear-
ing varieties (Lavrova et al., 1996; Korsakov et al., 2002). The calcsilicate rocks
are interlayered with garnet-biotite gneisses. Eclogites and amphibolites occur
as boudins in a matrix of gneisses and schists. A peculiar rock type of the Barchi
area are clinozoisite gneisses, which are often diamondiferous (Korsakov et al.,
2006, 2002).
In the Barchi Kol’ unit, rocks dip steeply to the north at 70o. In the central
part of the Barchi Kol’ UHP site three units can be destinguished (Fig. 2.2): the
northern unit, containing largely eclogites, the central unit, which has abundant
gneisses and the southern unit which is dominantly composed of granites and
orthogneisses. The lithologies are variable and it is hard to trace any particular
layer from one drill core to another (Korsakov et al., 2002). Masago (2000) studied
metabasites from the surface of Barchi area and mapped three zones with different
metamorphic conditions: in zone D rocks achieved UHP conditions estimated at
64
Figure 2.2: A) Simplified map of the Kokchetav complex based on the mapby Zhimulev (2007). B) Map of Barchi Kol’ area and samples location withsamples labels marked by red font. The geological map is compiled from mapsby Korsakov et al. (2002) and Masago (2000). Legend: 1 — UHP gneisses,garnet-pyroxene and silicate-carbonate rocks; amphibolites and micaschists; 2 —Eclogites; 3 — Grt-Qtz and Grt-Px rocks; 4 (a) — orthogneisses, (b) granites;5 — Medium grade metamorphic rocks. (a) zone C according Masago (2000),(b) - zone B; 6 — Low grade metamorphic rocks; (a) — Quartzites and schists(Kokchetav Formation); (b) — Mica schists and limestones (Sharyk Formation);7 — Pyroxenites and carbonatites (Barchi Massif, Krasnomai alkali-utrabasiccomplex).
65
27–40 kbar, 700–825oC; and in zones B and C peak metamorphic conditions
were much lower at 11.7± 0.5 kbar, 700± 30oC and 12–14 kbar, 700–815oC,
respectively (Fig. 2.2).
2.3 Analytical methods
Phase relations were analysed in polished thin sections using an optical micro-
scope and by back-scattered electron (BSE) images on a JEOL 6400 scanning
electron microscope (SEM) (Electron Microscopy Unit, ANU). The phase com-
positions were determined by EDS SEM, using an acceleration voltage of 15 kV, a
beam current of 1 nA and an acquisition time of 120 s. Distribution of major and
trace element in thin sections was mapped with a Cameca SX100 microprobe.
Fe, Mg, Mn, Y and P in garnet were measured using WDS spectrometers, with
Ca simultaneously analysed by EDS. The probe current and accelerating voltage
were 100 nA and 15 kV, respectively. The acquisition time for garnet maps was
from 2 to 12 hours. This permitted observation of zoning in garnet for P and
Y at the level of 1000 ppm. EDS analyses of mineral inclusions were monitored
contributions from the host mineral. Zoning of monazite was identified by high-
contrast backscatter electron (BSE) imaging using a Cambridge S360 scanning
electron microscope (SEM) at the ANU Electron Microscopy Unit (EMU; 2 nA,
15 kV and 15 mm working distance).
Trace elements in minerals were analysed by LA-ICP-MS at RSES, ANU,
using a pulsed 193 nm Ar-F Excimer laser with 100 mJ source energy at a rep-
etition rate of 5 Hz (Eggins et al., 1998) coupled to an Agilent 7500 quadrupole
ICP-MS. Laser sampling was performed in an He–Ar–H2 atmosphere using a spot
diameter of 25–37 μm. Data acquisition was performed by peak hopping in pulse
counting mode, acquiring individual intensity data for each element during each
mass spectrometer sweep. A total of 60 s, comprising a gas background of 20–25
s and 30-35 s signal, were acquired for each analysis. LA data were processed
by Excel spreadsheet created by Charlotte Allen. Trace element data in garnet
was calculated with NIST 612 (Pearce et al., 1997) as the external standard and
SiO2 as the internal standard. Monazite, rutile and zircon were calculated with
NIST 610 (Pearce et al., 1997) as the external standard and Ce, Ti and SiO2
as the internal standards, respectively. LA-ICP-MS of monazite with very low
HREE content demonstrated apparent positive anomalies of Er166 and Y b172 on
chondrite normalized patters. They were interpreted as interferences with oxides
of Nd150 and Gd156 or dioxide of Ce140, which are abundant in LREE-rich mon-
66
azite. Therefore Er and Yb were calculated from geometric averages of adjacent
REE normalized to chondrite. BCR-2 glass was employed as secondary standard
and its composition was reproduced within 10 %.
Raman spectra were obtained in Geoscience Australia with help of Terry
Mernagh. The Raman equipment comprises a Dilor SuperLabram spectrometer,
with a holographic notch filter (600 and 1800 g/mm gratings), liquid nitrogen-
cooled 2000 pixel CCD detector, and a 514.5 nm Melles Griot 543 argon ion laser
(5 mW at the sample). The spectral resolution was set at 2 cm−1 (slit width
of 100 μm). The microscope uses a 50X ULWD Olympus microscope objective,
focusing the laser spot to 2 μm in diameter and 5 μm deep.
U,Th-Pb isotope analyses of zircon and monazite were performed using the
sensitive, high-resolution ion microprobes at the RSES (SHRIMP II and RG)
using a 3.5–4.0 nA, 10 kV primary O−2 beam focused through a 120 μm aperture
to form a 25 μm diameter spot. Data acquisition followed Williams (1998) and
data were collected as sets of six scans throughout the mass range. The common
Pb correction was based on the measured 204Pb (Williams, 1998) and assum-
ing the Broken Hill common Pb composition. The measured 206Pb/238U ratios
were corrected using reference monazite 44069 (Aleinikoff et al., 2006, 425 Ma)
and TEMORA zircon (Black et al., 2003, 417 Ma). For monazite energy filter-
ing was used to eliminate interferences on 204Pb, as described by Rubatto et al.
(2001). Ages were calculated using Isoplot and SQUID software (Ludwig, 2003).
SHRIMP data were processed normalised to secondary beam monitor. Calibra-
tion errors were between 0.35 % and 2.4 % and calibration error was propagated
to individual analyses and averages. 1σ error is reported for individual measure-
ments and averages reported at 95% confidence level. If uncertainty of SHRIMP
dates was below 1 %, which is precision limit for studied range of ages, then
uncertainty was force to 1 % level.
2.4 Strategy
This study presents detailed and comprehensive comparison of the typical UHP
gneiss (B118A50) of the Kokchetav complex with somewhat different metased-
imentary samples (B94-333, B94-256 and B01-3). Samples B94-333, B94-256
and B01-3 have been selected based on their peculiar textural features, such as
homogeneous texture (B94-256, B01-3), large garnet porphyroblasts (B94-333),
and high abundance of mica (B94-256, B01-3). Sample B118A50 described in this
chapter is representative of typical diamond-bearing UHP garnet-biotite gneiss
67
of the Kokchetav complex and has the same texture and trace element compo-
sition as many other Grt-Bt gneisses. Also sample B94-26 described in detail in
Chapter 4 could be considered as another UHP gneiss suitable for a comparison
with other metasediments.
Three samples for this study (B94-333, B94-256 and B118A50) were selected
from the large collection of drill core material obtained during exploration of the
Barchi Kol’ area by the Kokchetav Prospecting Expedition. One sample (B01-3)
was collected from surface in the western part of the Barchi Kol’ lake area (Fig.
2.2). Numerous mapping campaigns of the Kokchetav complex demonstrated
that surface samples represent the underlying geology and are not transported
large distances. Therefore, sample B01-3 is considered likely to originate from
the UHP terrain.
In order to constrain metamorphic evolution of the samples and their relation
with UHP rocks they were systematically investigated for major and trace element
zoning of garnet, zircon and monazite, mineral inclusions in these minerals and
U-Pb ages were determined by SHRIMP dating of monazite and zircon. The
information obtained by these methods is summarized in Table 2.1.
2.5 Sample descriptions
B01− 3: is a weakly foliated, homogeneous, mica-rich schist of metapelitic com-
position (Fig. 2.3). The major mineral assemblage consists of Grt, Qtz, Ky, Phe,
Bt with accessory Rt, Mnz, Ap, Gr and Zrn. Phengite occurs as large euhedral
flakes forming the foliation of the rock. Grains of kyanite are small and often
have irregular, resorbed shapes and are enclosed in phengite. Garnet crystals are
euhedral and have a bimodal distribution: large crystals of ∼3 mm and smaller
grains <0.5 mm. Biotite is a minor constituent and it is associated with garnet
and phengite rims.
B94− 333: Garnet-biotite gneiss from drill hole. This rock contains Grt, Qtz,
Bt, Phe, Pl, Kfs, Ky and accessory Rt, Mnz, Zrn, Aln and sulfides. The rock has
a gneiss texture with thin layers composed of quartz-feldspar material and darker
layers enriched in biotite (Fig. 2.3). One such layer contains abundant grains of
rounded Ky, whereas, in other parts of the sample Ky is absent. Another layer
contains several large, elongated grains of pink garnet (up to 9 × 5 mm) which
has adjacent pressure shadows filled with quartz and feldspar. Another layer is
composed of Grt, Bt, Qtz and contains Aln. Garnet in this layer has an orange
colour.
68
Figure 2.3: Petrographic features of the studied samples on photographs, trans-mitted light and SEM images.
69
Figure 2.3 caption: a – e Sample
B01-3. a) Thin section of the sample B01-3 with large and small garnet grains
and large phengite flakes. b) garnet grains contain inclusions of graphite and
sometimes are in direct contact with monazite. c) Garnet porphyroblast with
LA-ICP-MS profile and e) enlargement of the core zone with inclusions of quartz
and xenotime. d) mantle zone of garnet with polyphase inclusion. (f – i) Sample
B94-333 f, g) scans of the sample. Striped texture with large garnet grains in
one part of the sample is visible. h) Large garnet with inclusions of rutile and
quartz in the mantle surrounded by the matrix of quartz, feldspars and biotite.
i) Apatite replacing monazite near large garnet crystal. j,k) Scan of sample and
thin section of sample B94-256, which show its foliated texture and association
of biotite with phengite and garnet. l) Inclusions of rutile and phengite-biotite
aggregate in garnet. Rutile is absent in matrix and its inclusions likely represent
high PT assemblage. Inclusion of mica aggregate shows gradual replacement of
phengite by biotite. (m–o) Sample B118A50. m) scan of sample and of thin
section (n). o) Photo of the thin section with inset showing zircon with diamond
inclusion. Note the difference in texture between samples B94-333 and B118A50
which have gneissic texture and samples B01-3 and B94-256 which have foliated
texture.
B94− 256: is a weakly foliated, homogeneous, mica-rich schist of metasedi-
mentary composition (Fig. 2.3). The sample is composed of Grt, Qtz, Kfs, Phe,
Bt and accessory Zrn, Mnz, Ap, Aln and sulfides. Garnet grains are approxi-
mately 2 mm in size. Phengite crystals are large (0.5–1 mm) and elongated along
the foliation. Biotite occurs as small grains along the edges of Phe and Grt and
has random orientation. Zircon and monazite form large grains (often >200μm)
in the matrix sample and monazite sometimes has a corona of apatite. Rutile is
present only as inclusions in garnet and is absent in matrix.
B118A50 is composed of Grt, Qtz, Kfs, Pl, Phe, Bt and accessory Zrn, Rt
and sulfides. The sample has a gneissose texture with thin quartz-feldspatic
layers and layers enriched in garnet and biotite (Fig. 2.3). Garnet grains are
small (∼2 mm), often fractured and surrounded by biotite and chlorite. Grains
of biotite and phengite have approximately the same size, shape and foliation.
Both biotite and phengite are significantly altered: biotite contains needles of
rutile and is partially replaced by chlorite, phengite is surrounded by chlorite
rims. Feldspars grains have an irregular shape and are also significantly altered.
70
The rock contains large grains of rutile aligned with the foliation, which are
partly replaced by mica and secondary rutile. Also aligned with the foliation
is aggregate of phengite and Th-REE minerals, which are pseudomorphs after
allanite.
2.6 Results
2.6.1 Mineral compositions and zoning
B01− 3:
Sample B01-3 garnet grains have extensive growth zoning in major and trace
elements (Fig. 2.4, 2.5, 2.6, 2.7, 2.8): large grains have a core with a uniform
content of MgO, CaO, FeO, MnO (Alm 83.5–86 %, Py 6.5–8 %, Grs 5–5.5 %, Sps
2.5–3.1 %), surrounded by a mantle with lower Mn content (Alm 81–86 %, Py 8–
13 %, Grs 4.6–5 %, Sps 1.2 %). In the rim, Mn and Fe decrease while Ca and Mg
increase (Alm 74–78 %, Py 13–17 %, Grs 7–8 %, Sps 0.2–0.9 %). The composition
of small Grt grains is identical to the composition of the rims of larger garnets.
HREE and Y concentrations from Grt core to rim decrease by a factor of 80
(Fig. 2.6). Garnet core and mantle REE patterns have a negative Eu anomaly
(Eu/Eu∗ = 0.1), which is reduced in the rims (Eu/Eu∗ = 0.5−−0.6). Garnets
show significant increase in P content from 40–50 ppm in the cores to ≈400 ppm
in rims (Fig. 2.6). Small garnet grains host up to 630 ppm P. Scandium contents
increase from 10 ppm in core to 90 ppm rim. Rutile grains in the matrix contains
250–370 ppm Zr. Large Phe flakes have cores with higher Si and Mg content,
while Ti increases from core to rim (Fig. 2.4, core: Si=3.28–3.32 pfu, 0.7–0.8
wt.% TiO2; rim: Si=3.18–3.21 pfu, 1 wt.% TiO2).
B94− 333:
Large pink garnet from sample B94-333 has three zones with different com-
positions. The garnet core is rich in MnO (Alm 73%, Py 14%, Grs 9 %, Sps 5%,
Fig. 2.5), the mantle-rim has lower Mn content but is rich in CaO and MgO
(Alm 56 %, Py 24 %, Grs 18 %, Sps 3 %). The HREE and Y contents decrease
from core to rim (Yb from 370 to 20 ppm; Y from 1400 to 160 ppm). The thin
rim differs from the mantle only by slightly larger negative Eu anomaly, whilst
they have the same major element compositions as the thick mantle. (Fig. 2.7).
Phosphorus content in the core is around 100 ppm while in the mantle and the
rim it increases to 500 ppm. High-Ca garnet has a large homogeneous core with
a thin rim which is depleted in HREE (Fig. 2.7). Phengite in the matrix of the
71
CCa Mn
Mg P
Mg Ti
Ca Mg
a b
c d
e f
g h
core
mantlerimB01-3
Grt
B01-3
Grt
Phe
Grt
B94-256PheBt
Mnz
Phe
Bt
Figure 2.4: Maps of garnet and phengite zoning. (a–d) Zoning of garnet fromsample B01-3. (e,f) Phengite zoning from the sample B01-3. (g,h) Garnet zoningfrom sample B94-256.
72
Figure 2.5: Compositional profiles across garnet showing major component zon-ing.
73
Figure 2.6: Trace element zoning in garnet.
Figure 2.7: Garnet REE patterns normalized to chondrite.
74
Figure 2.8: Composition of garnet and phengite found in the matrix and asinclusions in other minerals. Mnz inc – inclusions in monazite, Zrn inc – inclusionsin zircon, Grt inc – inclusions in garnet. Gray arrows demonstrate proposedevolution of phengite composition.
75
sample has low content of Si and significant variation of Ti content (Fig. 2.8).
Rutile grains in the matrix have rounded shape and contain 920–1100 ppm Zr .
B94− 256:
Main part of the garnet grains in sample B94-256 is homogeneous both in
major and trace elements and has low-Ca composition (Alm 61–62%, Py 26–29
%, Grs 7 %, Sps 3–4 %), HREE and Y are low (120–140 ppm), and with a small
negative Eu anomaly (Eu*/Eu = 0.64–0.47) (Fig. 2.6). However along rims,
cracks and near inclusions, the composition of garnet is different (Fig. 2.7, 2.5)
with lower Mg and higher Ca (Alm 62 %, Py 21–23 %, Grs 10–11 %, Sps 3–4
%) HREE and Y, and the Eu anomaly is more pronounced (Eu*/Eu = 0.25-
0.28). Phengite and biotite grains are relatively homogeneous in the matrix.
Phengite contains 3.11–3.16 Si pfu and 1.8–2.1 wt.% TiO2; retrograde matrix
biotite contains 2.71–2.81 Si pfu and 2–3 wt.% TiO2 (Fig. 2.8). Rutile is present
only as inclusions in garnet and contains 1030–1240 ppm Zr.
B118A50:
In sample B118A50 the distribution of the Ca, Fe, Mg and trace elements in
garnet is completely homogeneous and only in the rims is there a slight increase
in Mn (Fig. 2.5). Phengite in the matrix contains 3.15 Si pfu and 1.5 wt.% TiO2.
Biotite is Ti-rich and with Si 2.7 pfu. Feldspars are represented by almost pure
albite and K-feldspar. Large grains of rutile contain 790–920 ppm of Zr.
2.6.2 Mineral inclusions
Mineral inclusions were investigated in thin sections (in Grt) and in mineral sep-
arates (in Mnz and Zrn). The compositions of inclusions of garnet and phengite
is compared with matrix minerals in Fig. 2.8 and summarized in Table 2.1.
B01− 3:
Large garnet crystals have cores crowded with small inclusions of quartz and
occasionally xenotime (Fig. 2.3). There are small inclusions of Zrn, Ap, Rt
and Phe in garnet mantles, as well as polyphase inclusions with the following
associations: Xen + Qtz; Rt + Ilm; Rt + Chlor + Phe; Xen + Rt + Ilm + Chl
+ Kfs (Fig. 2.3). Garnet rims have no inclusions. Phengite inclusions in garnet
mantles contain 3.24–3.32 Si pfu.
Monazite contains inclusions of Ky, Grt, Phe, Zrn and Rt which mostly occur
near the core-rim boundary. Garnet inclusions in monazite have a composition
similar to that of the rims of large garnets (low Mn, high Mg and Ca). One
monazite core contains phengites with 3.08–3.11 Si pfu, which are lower than
76
most of white micas in the sample. Zircons from sample B01-3 contain inclusions
of Qtz, Grt and Phe which are too small for analysis.
B94− 333:
Large low-Ca garnet grains contain inclusions of rutile in the mantle (Zr
content in rutile inclusions is 880–1120 ppm, comparable to the matrix rutile) and
monazite inclusions close to the rim. Monazites contain inclusions of phengite
and garnet. Despite the high abundance of feldspar in the rock matrix, there
were no inclusions of feldspar in monazite. Phengite inclusions in monazite have
higher Si contain than matrix mica (Fig. 2.8). The single inclusion of garnet in
monazite is Mg rich (Fig. 2.8).
Zircons contain inclusions of Grt, Cpx and Phe. Phengite inclusions are simi-
lar to that included in monazite (Fig. 2.8). Garnet inclusions in zircon form two
separate groups: one group of inclusions resembles the cores of high-Ca garnet,
whereas the other inclusions are similar to the mantles-rims of the low-Ca garnets
(Fig. 2.8). The clinopyroxene is omphacite with 40 % of jadeite.
B94− 256:
Garnet contains inclusions of Rt, Bt, and Phe. Phengite inclusions in garnet
contain 3.15 Si pfu, 1.7 wt.% TiO2 and one inclusion in the garnet mantle con-
tains 3.28 Si pfu. Some inclusions of phengite have textures indicating partial
replacement by biotite, and thus inclusions of biotite are of secondary origin.
Rutile inclusions in garnet contain 1050–1250 ppm Zr, whereas rutile is absent
in the matrix. Monazite contains inclusions of Kfs, Bt and Phe. Inclusions of
phengite show a wide range of Si contents (3.1–3.36 Si pfu) with constantly low
Ti content (0.1–0.7 wt. %) (Fig. 2.8).
Zircons contain inclusions of Phe, Grt and Rt. Garnet inclusions in zircon
have a lower Fe content and higher Ca than the matrix garnet (Alm 58–60 %, Py
28–31 %, Grs 8–11 %, Sps 3.3 %; Fig. 2.8). Phengite inclusions in zircon have
high Ti (2.1 wt. % TiO2) and Si (3.3 Si pfu) contents, which differ from both
the matrix phengite and the inclusions in monazite.
B11A50:
Zircon from sample B118A50 contains inclusions of diamond and Grt, Cpx,
Phe and Bt. Garnet inclusions show a much more variable content of major
components than the matrix garnet (Fig. 2.8). Ti and Na contents in garnet
inclusions within zircon are similar as in matrix garnet or below detection limit
of EDS. Inclusions of Cpx in zircon contain 13–57 % of jadeite component and a
high Ca-Eskola component, whereas pyroxene is absent in the matrix. Phengite
inclusions in zircon have higher silica (3.2–3.4) and Ti (1.5–3 wt. % TiO2) content
77
than the matrix mica. Inclusions of metamorphic diamonds are present in CL-
bright mantle zones (which also have the highest Ti-in-zircon temperatures) and
at the boundaries between these zones and CL-dark cores.
2.6.3 Monazite and zircon description and geochronology
results
Monazite and zircon are resistant and robust minerals, and they preserve in-
dicators of high/ultrahigh pressure conditions. Also monazite and zircon are
chronometers and can provide information on the timing of metamorphic events.
Therefore internal zoning and trace element composition of monazite and zircon
were investigated. Separate growth zones of monazite in samples B01-3, B94-256
and B94-333 and zircon in samples B01-3, B94-256 and B118A50 were dated by
SHRIMP.
B01− 3 monazite:
Monazite from B01-3 is bright yellow and has a rounded shape. BSE images of
monazite reveal weak core-rim zonation (Fig. 2.9). Cores of grains usually show
a mosaic or polygonal-zoned texture and are brighter than rims, while the rims
are homogeneous on BSE images. Monazites have low Y and HREE contents (Lu
0.06–0.46 ppm, Y 100–800 ppm), which increase form cores to rims by a factor
of 8. Monazite cores in some grains have elevated strontium levels (Sr 1000–3000
ppm) decreasing to ≈800 ppm at the rim (Fig. 2.10). Th and Ca content also
decreases to the rims similar to Sr.
28 spot analysis on 14 grains were made by SHRIMP II; 4 spots were ex-
cluded because they were located near scratches and showed anomalously old
dates. 206Pb/238U dates scatter between 512±10 Ma and 537±10 Ma. 11 cores
give weighted average date of 529 Ma±7 Ma, the weighted average date of 13
rims is 525±7 Ma. Ages of cores and rims overlap and their averages are not
distinguishable at the precision of analyses. Average 206Pb/238U date for all 24
spots is 528±8 with MSWD of 1.5. The concordia age for the same data is 526±7
Ma (Fig. 2.11).
B01− 3 zircon:
Zircons are light pink, small (50-70 μm) and elongated (width:length ratio is
1:2). The internal CL zoning is complex: grains have small CL-dark cores with
weak sector zoning, surrounded by thin, CL-bright zones and outer CL-gray rims
with weak oscillatory-sector zoning. In transmitted light, many grains contain
yellow-coloured core domains with thin zoning (Fig. 2.12). However on CL
78
Figure 2.9: Back scattered electron images of monazite crystals. Circles marklocation of LA-ICP-MS and SHRIMP analyses and numbers show U-Pb dates inMa±1σ and Y content in wt.%. a) Monazite from micaschist B01-3 has core-rimzoning. b) Monazite from sample B94-333 has variable zoning. c) Monazite fromsample B94-256 often has zones with homogeneous BSE-zoning and others withpatchy zoning.
79
Figure 2.10: Trace element composition of monazite from studied samples andalso allanite in B94-256.
80
Figure 2.11: Results of SHRIMP U-Pb dating of monazite and zircon. Data-pointerror ellipses are 2σ. Sample B01-3 zircon ellipses are colored according Th/Uratio. Note, that analyses deviating from the main population have elevatedTh/U ratios. For other samples ellipses are colored according to their domain.
81
images these domains do not have specific zonation: they can be either CL-dark
or have thin, high-contrast zoning. BSE images of these cores do not show any
bright inclusions (at the resolution of Jeol 6400 BSE system). In zircons from
sample B01-3 only unspecified analyses (without distinction to cores or rims)
were made by both SHRIMP and LA-ICP-MS. The trace element patterns of the
zircons form two groups. (1) Zircon domains with normal patterns: enrichment
in HREE with respect to LREE, very low Th content (< 10 ppm) and 300–
700 ppm U (Fig. 2.13). These zircons have a slight concave pattern with a
depression at Ho, Er, Tm, Eu negative anomaly and positive Ce anomaly. (2)
Zircons with anomalous patterns, with a nearly flat REE distribution at 200–5000
times chondrite levels, strongly enriched in LREE with respect to a typical zircon
and rich in HREE, P (up to 4000 ppm), Ti, Y and Th. They have small negative
Eu anomalies and some of them show a small positive Ce anomaly. Grains with
this anomalous composition have the same Zr/Si ratio as a normal zircon. SEM
investigation of zircon zones with these unusual patterns did not reveal presence
of inclusions.
Zircons with normal REE patterns contain 2–10 ppm Ti, which corresponds
to T=645–720oC using the Ti-in-zircon thermometer of Ferry & Watson (2007).
Zircons with high-LREE patterns have elevated Ti concentrations (10-60 ppm),
which correspond to temperatures up to 930oC. However, these temperatures
were excluded because of anomalous composition of the zircons makes application
of any thermometer pointless.
15 grains were dated by SHRIMP II. 4 analysis are discordant, have high Th
content (1100–1700 ppm) and high Th/U ratios (1.8-2), and thus were excluded.
The remaining population still shows a significant spread of ages: from 504±7
Ma to 524±6 Ma. Spots deviating significantly from the main population have
elevated Th/U ratios, whereas those with a lower is Th/U ratios tend to fall
closer to the centre of the population (Fig. 2.11, ellipses are colored according
Th/U ratio). Average 206Pb/238U date corrected to 208Pb is 515±7 Ma. MSWD
is high 4.8 reflecting scatter.
B94− 333 monazite:
Monazite from sample B94-333 forms gray-yellow, transparent and relatively
rounded grains. Monazite occurs in the rock as inclusions in garnet, while in the
matrix it forms complex reaction aggregates (see section 3.6.1). Most monazite
grains from sample B94-333 are homogeneous in BSE images and a few grains
have thin oscillatory zoning in the cores (Fig. 2.9). Monazites show large range of
compositions with Lu varying 5–210 times chondritic levels, Y 0.02–1 wt.% (Fig.
82
Figure 2.12: Zoning of zircons in sample B01-3. (a) CL images of zircons revealscomplex zoning. (b) Corresponding transmitted light image showing zircons coreswhich have fine scale zoning and yelow coloration. Red ellipse denote the samegrains.
83
Figure 2.13: REE patterns of zircons.
1000
2 i5 100 c: 0
.s= ~ 10 c:
~ ·;:::;
0.1
--Normal pattern --Unusual pattern
801-3
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1000r------------------------,
2 :§
100
c: 10 0 .s= ~ ~ ·;:::;
0.1
O.Q1
--large, homogeneous - small, bright core
894-333
0.001 -1-~..---.--,--..----.--.---,--,.--,.--,---.--,..--,.--l
10000.------------------------,
1000
2 :g 100 c: 0
.s= 10 --core ~ --rim c: --inherited core 0 0 .... ·;:::;
0.1 894-256 O.D1
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
1000
100 2 :g
10 c: 0 .s= - core ~ - mantle c: - v_ery biright maotle 0 l::? - nm
·;:::; 0.1
0.01 811 8A50
84
2.10), Th 1.7 –11 wt.% and U 0.4–1.7 wt.%. The Sr content in monazite is high
(0.15–1 wt.%) and is not correlated with either monazite textures or the HREE
content. Ba is as low as 15-490 ppm. Monazite inclusion in garnet and individual
grain in the rock matrix were additionally analyzed in the thin section (Fig. 2.3).
Monazite inclusions in garnet have a significantly higher HREE content (Y 0.35
wt.%) than matrix monazite (Y 0.15 wt.%).
Monazite was dated by SHRIMP RG. The dates range between 515±8 Ma
and 538±8 Ma with an average 206Pb/238U date for 9 spots of 528±7 Ma, MSWD
1.1 (Fig. 2.11). No correlation between monazite age and composition or texture
was observed.
B94− 333 zircon:
Zircons from B94-333 occur as large pink grains with rough surfaces and wide
variety of shapes. There are two types of zircons according to their CL zoning and
size (Fig. 2.14). One zircon type forms small 50-100μm grains with CL-bright
cores and CL-dark rims. Another zircon type forms large (>200 μm) crystals
with dark almost structureless CL signal. Zircons of this type occasionally have
very thin rims with bright CL. The small zircons with CL-bright cores have
HREE enriched patterns, with significant variation in HREE content (10–120
times chondrite, Fig. 2.13) and low U content (<200 ppm). There is decrease
in the HREE content from CL-bright cores to CL-dark rims, whereas Ti content
is indistinguishable in two types of zircon. The large, CL-dark zircons have flat
HREE patterns with higher LREE contents, higher U concentrations (200–1200
ppm) and form a tight cluster of compositions. Th/U ratios are low (0.01–0.15)
in both populations. The Eu anomaly is similar in both zircon types (Eu∗ 0.6–1)
and Ti concentrations are 20–52 ppm, which correspond to T 815–940oC (Ferry
& Watson, 2007) (Fig. 2.15). Zircon from sample B94-333 were not dated.
B94− 256 monazite:
In sample B94-256 monazite grains have a yellow-greenish colour and rounded
or tabular shape. BSE images reveal that the large grains have two zones: one
zone is devoid of zoning and the other was a patchy texture (Fig. 2.9). Small
grains appear either homogeneous in BSE or have thin bright rims on darker
cores. Texturally the monazite with patchy zoning formed after the homogeneous
monazite because the patchy monazite domains form the outer zones of grains
and overprints the homogeneous zones (Fig. 2.9).
Homogeneous monazite zones have higher HREE, Y, Ba (about 300 ppm) and
Sr (500–1500 ppm) contents than the patchy monazite (Ba < 200ppm, mostly Sr
< 320 ppm, Fig. 2.10). Monazite domains with patchy zoning show a constant
85
Figure 2.14: Zircon internal zoning revealed by CL imaging. Circles show spotsof LA-ICP-MS and SHRIMP analyses (Ma±1σ) and numbers show U-Pb datesin Ma
86
Figure 2.15: Histogram of Ti-in-zircon temperatures calculated by thermometerby Ferry & Watson (2007) for various zircon types and growth zones.
87
Th content (2.1–2.9 wt.%) and variable U content with Th/U ratios between
3.4–7.2. Monazite with homogeneous zoning has variable Th (0.8–2.6 wt.%) with
Th/U = 3–6, except one analysis where Th/U=10. Mineral inclusions mostly
occur in patchy zones.
21 U-Pb analyses on 14 grains were made by SHRIMP II. One analysis was
excluded because of high 204Pb. 206Pb/238U dates range from 510±6 Ma to 533±6
Ma. Ages of homogeneous and patchy zones in the same grains are within error.
Average dates of patchy and homogeneous zones overlap and are not distinguish-
able at the precision of the analyses (Fig. 2.11). Average date for all spots is
521±6 Ma. All spots are concordant and the concordia age for all the spots is
521±6 Ma (Fig. 2.11 b). Dating of 9 homogeneous domains yield a date of 515
Ma±7 Ma and 12 patchy zones date of 523±5 Ma.
B94− 256 zircon:
In sample B94-256 the zircon crystals are large, clear, light pink, round and
with a smooth unfaceted surface. CL images of these zircons reveal very consis-
tent core-rim zoning with fine scale oscillatory zoning, or CL-dark cores and thick
CL-bright rims (Fig. 2.14). Rims have diffuse areas of very bright CL. These
bright areas are associated with cracks and edges of crystals and probably caused
by strain in crystals. Similar textures have been described in zircons from HP
rocks (Corfu et al., 2003).
Both zircon cores and rims have REE patterns with flat HREE, LREE deple-
tion, a small negative Eu anomaly, and a positive Ce anomaly. Cores have HREE
contents 150–250 times chondrite values. One core has high LREE content and
increasing HREE pattern. Rims have higher REE content equal to 300–400 times
chondritic values (Fig. 2.13). Ti concentrations range from 12–84 ppm in the
cores to 100–150 ppm in the rims, for which Ti-in-zircon thermometer provides
temperatures of 765–990oC and 1020–1080oC, respectively (Ferry & Watson,
2007) (Fig. 2.15). There is a positive correlation between Ti concentrations and
HREE and Y concentrations and Th/U ratios.
15 U-Pb analysis on 8 grains were measured by SHRIMP II. Spot 206Pb/238U
dates (208Pb corrected) span from 516±9 to 530±8 Ma (Fig. 2.11), and one spot
in a core with high content of LREE and high Th/U ratio gave an old 207Pb/206Pb
date of 2867±70 Ma. There is no systematic difference between the ages of the
cores and the rims (core 522±6 Ma, and rims 522±7 Ma). The weighted average
is 522± 6 Ma (MSWD=0.6, N=14).
B118A50 zircon:
Sample B118A50 was strongly depleted in LREE by UHP melting (Chapter
88
3), and allanite most likely formed during exhumation. Allanite decomposed on
retrogression to Th-minerals and monazite.
In sample B118A50 zircons are isometric, almost colorless, with a euhedral,
soccer-ball shape. Zircons from this sample show concentric zoning and several
distinct CL zones can be specified: CL-dark cores, CL-bright mantles, mantles
with very bright CL and CL-dark rims. The CL-dark cores have low Th/U
ratios <0.05 and 9–63 ppm Ti (Ti-in-zircon temperatures of 850–950oC). These
cores are overgrown by CL-bright mantles. These zones have 46–108 ppm Ti,
high Ti-in-zircon temperatures (910–1030oC) and Th/U ratios tend to be higher
(0.04–0.3). Additionally some grains contain some very CL-bright mantle zones.
These zones have very low content of U (<100 ppm) and also high Ti content of
60–116 ppm (Ti-in-zircon temperatures 950–1040oC). CL-dark rims are present
on almost all zircons. These zones have Ti content of 32–64 ppm similar to the
cores (Ti-in-zircon temperatures 850–930oC), but with higher Th/U ratios (0.05–
0.15). Zircons from sample B118A50 have REE patterns with a flat distribution of
HREE, a small negative Eu anomaly and a large positive Ce anomaly. Variations
in trace element composition in zircons from B118A50 are relatively small. HREE
content varies from 30 to 100 times chondrite levels, and there is no variation in
the slope of these patterns.
18 U-Pb analysis on 12 grains were measured by SHRIMP II. One spot pro-
vided a discordant 206Pb/238U date of 617±9 Ma. This analysis was affected by
inheritance and thus was excluded. Another grain was excluded due to its very
low U content (<20 ppm) and the resulting large uncertainty. The 16 remaining
spots vary from 503±7 Ma to 532±7 Ma and the average is 520±7 Ma. The
high MSWD of 4.3 reflects scatter above analytical uncertainty. However, U-Pb
dates of cores, mantles and rims are overlapping and indistinguishable within the
precision of measurements (Fig. 2.11).
2.7 Discussion
This study presents new U-Pb data on monazite and zircon from HP/UHP
metasedimentary rocks from the Barchi Kol’ area of the Kokchetav complex. The
studied samples have large variations in minerals compositions, zoning and other
petrologic characteristics, which may be related to the variations in PT paths.
In relation to this study the most important information about HP metamorphic
conditions was obtained from Ti-in-zircon thermometer, garnet zoning and min-
eral inclusions. This section evaluates these methods and then presents estimates
89
of PT evolution. Trace element compositions of minerals, U-Pb geochronology
and PT estimates are combined in order to reach tectonic interpretations of the
studied samples. Finally, a comparison with other UHP complexes is presented.
2.7.1 Mineral inclusions in the Kokchetav rocks
Numerous studies demonstrated that zircon is the best host mineral for HP/UHP
inclusions because it grows significantly at those conditions but, does not react
with silicates, and has a very robust lattice (Sobolev et al., 1991; Hermann et al.,
2001; Katayama & Maruyama, 2009; Liu et al., 2001, 2002). The capacity of
monazite for preservation of HP-UHP minerals is less known though there is
report of diamond inclusion in monazite from quartzofeldspathic of Erzgebirge.
Potentially monazite is a good host fro inclusions, because as a REE phosphate
monazite has limited ability for cation exchange with silicate minerals. Zircon
and monazite from the studied samples contain inclusions of Phe, Grt, Cpx and
a SiO2 phase.
Inclusions of SiO2 in zircons were investigated by Raman spectroscopy. Ra-
man confirmed that zircon from sample B118A50 contains coesite and diamond
inclusions, which are typical for UHP gneisses. In sample B01-3 garnet contains
inclusions of quartz, which have shift of Raman peaks from overpressure (Kor-
sakov et al., 2009). Zircon from samples B94-256 and B94-333 were investigated
by Raman, but as yet neither coesite nor quartz have been unidentified in that
samples.
Phengite is an abundant inclusion mineral and has large variations in com-
position. The Si content in phengite increases with increasing pressure and de-
creases with increasing temperature, and Ti content increases with increases of P
and T (Hermann, 2002b; Auzanneau et al., 2010). Phengite inclusions in zircon
from sample B118A50 have the high Ti and Si contents (Fig. 2.8) in agreement
with a UHP origin. Inclusions of phengite in zircon from sample B94-256 has
high Si content, comparable with mica from UHP sample B118A50 and micas
produced in experiments at 35–45 kbar and 900–1000oC (Hermann & Spandler,
2008; Auzanneau et al., 2010). Phengite in sample B94-256 have very high Ti
content, up to 2.15 wt.% or 0.22 Ti pfu (calculated for 11 oxygens, Fig. 2.8).
In fact according to a compilation of experimental HP phengites by Auzanneau
et al. (2010) maximum Ti content is typically <0.15 Ti pfu and phengite with
higher Ti content of 3.5 wt.% was produced in the only one experiment at 850oC
23 kbar. This difference could originate from the over-saturation of experimental
90
compositions by water, and thus low maximum temperature of phengite stability.
Phengite inclusions in monazite and zircon from sample B94-333 have substan-
tially lower maximum Ti and Si content than in samples B118A50 and B94-256,
thus demonstrating lower peak conditions PT conditions. Mica in sample B01-3
has comparable maximum Si content but lower Ti content than sample B94-333,
and considering positive slope of lines of constant Si content in mica, sample
B01-3 had lowest peak PT among the studied samples.
2.7.2 Application of Ti-in-zircon thermometers to UHP
rocks
Solubility of Ti in zircon potentially is an attractive thermometer for UHP rocks
because zircon is reactive at high temperature and Ti diffuses exceptionally slowly
in zircon (Cherniak & Watson, 2007), when compared to the majority of other
minerals. Additionally, the Kokchetav UHP gneisses are favourable for the appli-
cation of the Ti-in-zircon thermometer because rutile and SiO2 phase were stable
during the entire PT evolution. Therefore, activities of SiO2 and TiO2 (aSiO2
and aT iO2) will be close to unity, and one of the major uncertainties associated
with this thermometer is constrained. Furthermore, zircon is known as a perfect
container for HP and UHP minerals — often HP-UHP minerals are found only
as inclusions in zircon, whereas in the rest of a rock they are completely retro-
gressed (Sobolev et al., 1991; Liu et al., 2001; Katayama & Maruyama, 2009).
These properties of zircon can help to obtain pressure estimates in addition to
temperatures. Ti-in-zircon thermometer was applied to HP/UHP rocks from
Dabie-Sulu complex, China (Gao et al., 2011; Zheng et al., 2011a) and demon-
strated different growth zones with different ages can be determined from Ti
content.
Unfortunately, the effect of pressure on the Ti-in-zircon thermometer has not
yet been calibrated experimentally. Watson et al. (2006) compared Ti concentra-
tions in zircons from the experiments performed at 10 kbar with natural zircons
from the Bishop tuff (which supposedly crystallised at 3 kbar) and the Labait
harzburgite (≈ 30 kbar) and concluded that the pressure effect is negligible on
the Ti solubility in zircon. Later, Ferry & Watson (2007) estimated the pressure
effect on the Ti-in-zircon to be 50oC/10 kbar. However, Tailby et al. (2011)
argued that aT iO2 and aSiO2 in the Bishop tuff and the Labait harzburgite were
unconstrained, thus making them unsuitable for calibration of the pressure effect
on Ti solubility. Tailby et al. (2011) performed experiments at controlled aT iO2
91
and aSiO2 at atmospheric pressure. The Ti solubility in zircon was more than
two times higher than predicted by the calibration proposed by Ferry & Wat-
son (2007). Hence Tailby et al. (2011) concluded that pressure has a significant
negative effect on the Ti solubility in zircon. Therefore it is expected that the
Ti-in-zircon thermometer should underestimate temperatures at high pressure
conditions.
Ti concentrations in zircons from the studied samples are very high (Fig.
2.15). For zircons from samples B94-256 and B94-118A50 temperatures calcu-
lated using the thermometer calibration by Ferry & Watson (2007) reach 1050-
1080oC. If an error in Ti concentration of 15 % is assigned, then the uncertainty
on the temperature is ±20 oC. Therefore, without any pressure correction, the
temperature estimated by the Ti-in-zircon thermometer (Ferry & Watson, 2007)
is close of higher than peak conditions of 950-1000 oC, >45 kbar estimated for
the Kokchetav rocks in other studies (e.g. Sobolev & Shatsky, 1990; Hermann
et al., 2001). If Ti-in-zircon thermometer underestimate temperature at high
pressure (Tailby et al., 2011) then peak metamorphic temperatures were sig-
nificantly higher than 1000oC. Conversely the pressure effect may indeed be
insignificant, or non-linear with increasing pressure. Considering these uncer-
tainties Ti-in-zircon will be used below in relative sense: similar temperatures
are considered representing similar PT conditions.
Application of the Ti-in-zircon thermometer shows that different samples
achieved different peak conditions. In UHP gneiss B118A50, zircons mantle zones
with diamond inclusions record Ti-in-zircon temperatures of 910–1040oC, thus
these zones correspond to peak conditions of UHP metamorphism (Fig. 2.15).
Sample B94-256 demonstrated Ti-in-zircon temperatures 990–1080oC , similar
to sample B118A50, what can be evidence for similar peak conditions in both
samples. However samples B01-3 and B94-333 demonstrated markedly lower
maximum Ti-in-zircon temperates (645–720 and 815-940oC respectively) and this
indicates lower peak conditions in these rocks. Alternative hypothesis is that zir-
con formed not at peak conditions. However at peak conditions the Kokchetav
rocks melted, and at the presence of melt always there is new zircon growth.
2.7.3 Interpretation of garnet zoning
Samples demonstrate remarkably different types of garnet zoning in major and
trace elements (Fig. 2.5, 2.6). Garnets in sample B118A50 are essentially homo-
geneous, and in sample B94-256 garnets record only late retrogression. Contrast-
92
ingly in samples B94-333 and B01-3 garnet have high concentrations of HREE
and Y in cores, which decrease to the rims by orders of magnitude. Intensive
zoning in HREE content is typical for garnets from medium grade metapelites
(Yang & Pattison, 2006; Spear & Pyle, 2010), because HREE and Y content in
garnet are controlled by the presence of xenotime, and equilibrium between gar-
net and monazite. When xenotime is completely dissolved then garnet becomes
the main host for HREE and Y. At high temperature HREE and Y start to
diffuse in garnet with sufficiently high speed to homogenise the grains. Because
diffusion of REE in garnet is significantly slower that of major elements REE
zoning is erased by diffusion only at significantly higher temperature than major
elements zoning (Van Orman et al., 2002; Hermann & Rubatto, 2003). Sample
B01-3 preserves high concentrations of Mn in the core. Mn is an element that
is very compatible in garnet, and hence is typically sequestered in metapelites
by garnet on the early stages of growth. Hence garnet zoning allows samples
to be arranged according their peak temperatures: samples B118A50 and B94-
256 achieved highest temperatures above closure temperature for REE diffusion
in garnet, sample B94-333 achieved conditions above the closure temperature for
major divalent cations, but below REE closure temperature, and in sample B01-3
temperature was not sufficient even to homogenize major elements. These rela-
tive estimates of peak temperatures are in complete agreement with Ti-in-zircon
temperates and compositions of mineral inclusions.
Sobolev et al. (2011) reported the presence of zoning in oxygen isotopic com-
position in garnets from diamond-bearing marbles from the Kokchetav. Presence
of zoning in oxygen isotopes in UHP samples whereas trivalent rare earth ele-
ments in this study are found to be homogeneously distributed can be explained
by slower diffusion of oxygen in garnet 1, or late formation of the isotopic hetero-
geneity in garnets studied by Sobolev et al. (2011).
Garnet trace element composition potentially is an important tool for dis-
crimination of rocks with different PT paths. The presence or absence of an
Eu anomaly can be an indicator of the presence or absence of feldspar in the
assemblage. Along their PT evolution HP/UHP gneisses have passed though
HP assemblages where plagioclase should react completely to form omphacite
pyroxene, and then on retrogression pyroxene decomposed and new plagioclase
appeared. However in my samples, the majority of garnets and zircon analyses
show very small Eu anomaly comparable with small Eu anomalies typical for up-
1Diffusion of oxygen in garnet is very poorly constrained experimentally: the only referenceis to a PhD thesis with unknown experimental methods.
93
per crustal rocks and metasediments. One explanation for such behaviour can be
if retrogression of omphacite occurred with little involvement of garnet. Another
option is that in the Kokchetav rocks Eu was predominantly in Eu3+ state and
thus it was not fractionated by feldspars.
Ti and Zr can have temperature and pressure dependence in garnet (Hermann
& Spandler, 2008; Rubatto & Hermann, 2007a). Experimental studies reported
0.4–0.9 wt. % TiO2 and 200–350 ppm of Zr in garnets. In sample B01-3 Zr con-
tent is ∼5 ppm Zr, and in other samples it is 10–20 ppm Zr. Ti content increases
from ∼50 ppm Zr in sample B01-3, to 100–1000 ppm in B94-333, 110–240 ppm
in B94-256 and 500–1000 ppm in B118A50. Hence Ti and Zr concentrations are
much lower than in experiments. It can be either due to rapid loss of Ti and Zr
from garnet by diffusion, or different behaviour than in experiments.
2.7.4 PT paths
All investigated samples experienced prolonged metamorphic evolutions, where
mineral associations changed significantly at different conditions. Therefore in
interpretation of PT paths special emphasis was put on mineral inclusions in ro-
bust minerals and features which are least affected by retrogression such as garnet
trace element zoning, and trace element composition of zircon. The summary of
petrographic and geochronologic features for all samples is presented in Table
2.1, and reconstruction of the mineral evolution of samples is shown in mineral
evolution diagrams in Fig. 2.16.
Sample B01-3
Garnet in the sample shows prograde zoning with almost continuous decrease of
Mn and HREE from core to rim (Fig. 2.4, 2.5, 2.6). The PT-path for sample
B01-3 is calculated by applying a number of geothermobarometers:
# Garnet cores with inclusions of xenotime and quartz contain 1200–1400
ppm Y. Based on the temperature dependence of Y content in garnet in equilib-
rium with xenotime this gives a temperature of 510±5oC (Pyle & Spear, 2000).
The shift in the Raman spectra of quartz inclusions in garnet cores yields pres-
sures of ∼10 kbar (Korsakov et al., 2009). Two phengite inclusions in monazite
have low Si content, and differ from all other mica in the sample, thus presumably
originating from prograde part of PT path. Pressure estimates by the Grt-Phe-
Ky barometer (Ravna & Terry, 2004) from the garnet core composition and that
low-Si phengite inclusions in monazite give pressure estimate of 11 ± 2 kbar,
94
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Figure 2.16: Mineral evolution diagrams for the studied samples. Different thick-ness of lines demonstrates changes in mineral abundance. o — denotes homog-enization of mineral by diffusion. M inc — monazite inclusions, Z inc — zirconinclusions, mtx — matrix.
which is identical to estimate from quartz inclusions. Application to the same
compositions of Fe-Mg exchange geothermometer by (Green & Hellman, 1982,
low Ca, low Mg) yields temperature of 500oC.
# Garnet mantles contain xenotime and contain 600–400 ppm of Y, which
gives temperatures of 550± 10oC (Pyle & Spear, 2000). Calculation by the Grt-
Phe-Ky geobarometer (Ravna & Terry, 2004) for composition of garnet mantles
and of phengite inclusions in garnet mantle gives 23 ± 2 kbar. Fe-Mg exchange
geothermometer by (Green & Hellman, 1982, low Ca, low Mg) yields 570oC.
# Zircons contain from 2 to 10 ppm Ti, which corresponds to T=640–730oC
(Ferry & Watson, 2007). The sinusoidal shape of the REE pattern of zircons re-
95
flects that of the garnet rims, thus providing evidence that the zircons crystallised
together with the garnet rims. In the matrix, phengite crystals from core to rim
show a decrease in Mg (which is a proxy for Si) and an increase in Ti content.
This zoning can be formed only by increasing temperature, which increases Ti
solubility in phengite and a decrease in pressure, which controls Si in phengite
(Auzanneau et al., 2010). Rutile in the matrix contains 240–370 ppm Zr with
an average of 300 ppm. Simultaneous application of the Grt-Phe-Ky barometer
(Ravna & Terry, 2004) with the Zr-in-rutile thermometer (Tomkins et al., 2007)
yields a temperature of 690–730oC and pressure of 24 ± 2 kbar. Application to
the same compositions of Fe-Mg exchange geothermometer by (Green & Hell-
man, 1982, low Ca, low Mg) provided temperature estimate of 780oC. These are
the highest PT conditions recorded by samples B01-3 and thus represent peak
metamorphism.
These calculations constrain three points of the PT path for sample B01-3
(Fig. 2.17).
Sample B94-333
Several features indicate that sample B94-333 attained lower peak metamorphic
conditions than samples B92-256 and B118A50:
# Garnet cores have very high concentrations of HREE and REE zoning is
similar to zoning in sample B01-3. Therefore it can be interpreted as prograde
REE zoning. The temperature of core formation can be estimated, using the
B01-3 garnet as a proxy, by Y-in-garnet thermometer at ≈550oC (Pyle & Spear,
2000; Spear & Pyle, 2010). Such high concentrations of HREE in garnet were
not observed in garnets from other UHP gneisses from this study, presumably
due to diffusional homogenization. Therefore it is suggested that the closure
temperature for diffusion of REE in garnet was not attained in sample B94-333.
# Ti-in-zircon temperatures are 815–950oC with an average of 900oC. Though
the significance of absolute values of the Ti-in-zircon thermometer is unclear (see
2.7.2), Ti concentrations in zircons from B94-333 are clearly lower than in UHP
gneisses B118A50 and B94-26 thus demonstrating lower peak PT conditions.
# Omphacite inclusions in zircon indicates that the rock reached eclogite
facies. Phengite inclusions in zircon and monazite have higher Si content than
matrix phengite and similar Ti contents (Fig. 2.8), thus representing peak condi-
tions. Micas with compositions similar to inclusions in monazite and zircon were
synthesised in metasedimentary bulk composition at 800-850oC and 25-35 kbar
96
Figure 2.17: Proposed PT paths of studied samples. Ellipses represent uncer-tainty on PT estimates. The PT path of B118A50 is based on Hermann et al.(2001) with changes according to Auzanneau et al. (2006). Gray lines show loca-tion of reactions of phengite breakdown, with phengite disappearance right/belowlines: (a) reaction Phe(Ms)+Qtz(Coe)=melt from (Auzanneau et al., 2006), (b)reaction Cpx+Phe+Qtz=Bt+Pl+Grt+melt (Auzanneau et al., 2006), (c) phen-gite upper stability limit (Hermann & Spandler, 2008).
(Auzanneau et al., 2010).
Sample B94-333 presents some evidence of partial melting: the sample has
gneissic texture typical for migmaties (Sawyer, 2008), including Kokchetav UHP
gneisses, additionally some monazite grains have oscillatory zoning typical for
monazite grown from melt (Crowley et al., 2008). Phengite in matrix forms
small grains and has low Si and Ti content, probably due to breakdown of high
pressure phengite by melting. There are two types of garnet in sample B94-333:
pink low-Ca Grt, whose cores preserve prograde REE zoning, and orange, high-
Ca Grt. This variation in garnet composition suggests that this sample contains
layers with different bulk composition.
The Grt-Phe thermometer (Green & Hellman, 1982, low Ca, low Mg) for
the composition of matrix phengite and low-Ca garnet rims gives a temperature
estimate of 740±30oC. Application of the Grt-Pl-Ms-Qtz geobarometer (Hodges
& Crowley, 1985) to a matrix assemblage and the low-Ca garnet gives pressure
97
estimates of 12±3 kbars. The Zr in rutile thermometer (Tomkins et al., 2007)
applied to rutile inclusions in the garnet mantle and matrix grains gives 750–
770oC at 10 kbar and 800–815oC at 20 kbar (Tomkins et al., 2007). These data
constrain the PT path of sample B94-333 as drown in Fig. 2.17.
Sample B94-256
Several features of sample B94-256 suggest that it reached high P and T condi-
tions:
# Very high Ti-in-zircon temperatures (up to 1080oC, Fig. 2.15).
# High Ti and Si content in matrix phengite and inclusions in zircon (Fig.
2.8).
# Prograde zoning in garnet was completely erased by diffusion.
All these features in sample B94-256 are similar to the typical UHP gneiss
B118A50, thus indicating achievement of UHP conditions.
However, several features indicate that, contrary to other UHP gneisses, sam-
ple B94-256 did not undergo partial melting:
* Sample is homogeneous on the macroscale, distinguishing it from gneissose
textures of other UHP gneisses (compare Fig. 3.3 and 2.3 in Chapter 3).
* Most of the phengite inclusions in garnet have compositions identical to
matrix phengite, show textures of phengite replacement by biotite (Fig. 2.3) and
have halos of high-Ca garnet around them (Fig. 2.4). Thus, mica inclusions in
garnet were reset/reequilibrated with the matrix assemblage. Monazite contains
inclusions of K-feldspar and phengite with low Ti contents (0.1–0.7 wt.% TiO2)
much lower than matrix phengite (1.8–2.1 wt.% TiO2 and 3.1 Si pfu) and a wide
variation in Si content (3.15–3.35 pfu). As Ti in phengite is strongly temperature
dependent (Hermann & Spandler, 2008; Auzanneau et al., 2010), it suggests that
the inclusions in monazite represent a low temperature stage of the metamorphic
path with increase of pressure (e.g. the prograde part of PT path). On the other
hand, preservation of this inclusions implies that monazite was not recrystallised
during peak conditions or retrogression. However, in other UHP gneisses melting
at peak conditions caused complete dissolution of monazite in melt even at low
melting degree (Chapter 3). Phengite inclusions in monazite have very low Ti
content, lower than any mica from sample B01-3 thus its temperature of formation
can be estimated <500oC, Si content is high and pressure can be estimated as
≈20 kbar.
* Zircons from sample B94-256 have very low Th content and low Th/U ratios
98
(<0.04), that is explained by presence of monazite in the assemblages. Exception
is one inherited core with high Th/U ratio. In contrast, high temperature zircons
from samples B118A50 and B94-26 (Chapter 4) have high Th/U ratios because
at these conditions monazite was dissolved in the melt. In other UHP sam-
ples (B118A50 this Chapter and sample B94-26 in Chapter 4), zircon rims show
decreases in Ti-in-zircon temperatures because they formed during melt crystal-
lization. In contrast, Ti-in-zircon temperatures in sample B94-256 increase from
core to rim, and there are no outer zones with lower temperatures thus indicating
absence of melt in the sample during the cooling stage.
* Presence of melt is considered as one of main factors of poor preservation
of HP associations in the UHP gneisses (Hermann et al., 2001). In the matrix
of sample B94-256, phengite has a higher Ti content (i. e. higher temperature)
than in matrix of B118A50. Preservation of high T phengite probably indicates
subsolidus conditions in the B94-256, whereas in UHP gneisses B118A50 melt
caused retrogression of phengite.
Therefore, the sum of evidences indicates in B94-256 similar peak conditions
to other UHP gneisses, but the rock differs markedly by absence of melting.
Absence of melting in the sample can be explained by combination of two fac-
tors: low water activity and lower temperature below phengite melting during
exhumation than in other UHP gneisses.
Phengite inclusions in monazite show large range of Si content at low Ti
content with some tendency of higher Ti concentrations in low Si micas. Lines
of constant of Ti content in phengite (isoplets) have small positive slope with
pressure (Hermann, 2002b; Auzanneau et al., 2010), and thus the variation of
compositions of phengite inclusions can indicate PT path parallel to such isoplet
at low T conditions. Such evolution of mica inclusions can be interpreted as
subduction of the rock along a cold geotherm. The phengite inclusion in zircon
show much higher Ti content at comparable Si content. This can be summarized
in a kinked PT path that has a sharp increase of temperature around 30 kbar
(Fig. 2.17).
Rutile is present only as inclusions in garnet and Zr content of rutile corre-
sponds to a temperatures of 760-780oC at 10 kbar and 810–830oC at 20 kbar
(Tomkins et al., 2007). This temperature is significantly lower than peak tem-
perature for UHP gneisses thus likely Zr in rutile records late equilibration. In
the matrix rutile is absent probably due its consumption by biotite, which has
high Ti content.
In B94-256 the matrix assemblage is Grt-Qtz-Phe-Bt-Kfs. Biotite appeared by
99
replacement of garnet and rutile and its formation resulted in appearance of high-
Ca garnet with increased content of HREE and stronger Eu anomaly. Calculation
by the Grt-Phe thermometer (Green & Hellman, 1982, low Ca, low Mg) gives
temperatures of 700±20oC for garnet rims and matrix phengite. Application of
Grt-Phe geobarometer Ravna & Terry (2004) for this temperature gives pressure
estimate of 21±2 kbar (but Ky is absent in the association). These data allow to
constrained PT path of sample B94-256 (Fig. 2.17).
Sample B118A50
Sample B118A50 contains diamond inclusions in zircons (Fig. 2.3) and thus is
typical of the Kokchetav UHP gneiss. Important features of the sample include:
(1) very homogeneous garnet (Fig. 2.5, 2.6), which is evidence for high tempera-
ture diffusional homogenization, (2) high Ti-in-zircon temperatures (Fig. 2.15),
(3) specific trace element zoning of zircon, recording melting at peak conditions,
(4) high pressure mineral inclusions in zircon. On the other hand, similar to other
UHP gneisses, sample B118A50 has intense retrogression: (1) no omphacite is
preserved in matrix, (2) garnet and phengite are partly replaced by biotite, (3)
garnet and biotite are significantly retrogressed to chlorite, and (4) rutile shows
Zr-in-rutile temperatures of ≈800oC, much lower than the peak temperature es-
timate. It is assumed that sample B118A50 records the same peak conditions
and PT path as other UHP samples from Kokchetav constrained by previous
studies (Sobolev & Shatsky, 1990; Hermann et al., 2001; Chopin, 2003): peak
at 950-1000oC and >45 kbar, 800oC and 10 kbar, 550-600oC and 5 kbar (Fig.
2.17).
2.7.5 Linking accessory minerals with metamorphic evo-
lution
B01-3
B01− 3 Monazite:
Monazite from sample B01-3 contains inclusions of Grt, Phe, Ky, Rt and Qtz
– an assemblage identical to that of the matrix. Compositions of garnet and
phengite inclusions in monazites are also identical to those of garnet rims and
matrix phengite, respectively (Fig. 2.8), indicating that the growth of monazite
was contemporaneous with garnet rims. In garnet there is a large decrease in
HREE content from core to rim, while in monazite the HREE content is very
100
low and increases from core to rim. Dmnz/grtLu =0.1 between average monazite core
and average garnet rim, and Dmnz/grtLu =0.3 between monazite rim and garnet rim
are lower than values of Dmnz/grtLu =2–3 reported for the equilibrium monazite-
garnet partitioning in granulites (Hermann & Rubatto, 2003; Rubatto et al.,
2006). This difference in Dmnz/grtLu may be due the difference in conditions in the
granulites studied by Hermann & Rubatto (2003) and Rubatto et al. (2006) where
conditions were around 800oC and 4–8 kbar pressure. Increases of HREE content
in monazite rims can be explained by an increase of Dmnz/grt with temperature
as proposed by Pyle et al. (2001). Therefore, the trace element composition of
monazite provides additional evidence for its growth in equilibrium with garnet
rims. A possible precursor for monazite may be earlier monazite, allanite, REE
carbonates or REE-bearing Al-phosphates (Yang & Pattison, 2006; Janots et al.,
2006).
B01− 3 Zircon:
In sample B01-3 zircons have domains with “normal” and “anomalous” REE
patterns. Zircons with “normal” patterns have sinusoidal REE patterns with
small depressions in Ho, Er, Tm relative to Tb, Dy, Yb, Lu. This feature can be
explained only by zircon’s equilibrium with garnet rims, which show a concave
REE pattern (Fig. 2.13, 2.7). Core and mantle zones of garnets have REE
patterns with a rising or flat HREE distribution and cannot be in equilibrium
with zircons. Dzrn/grt for HREE slightly increases from 0.8 for Gd-Ho to 1.3 for
Lu. These values are significantly lower than Dzrn/grt of 6-8 found in granulites
(Hermann & Rubatto, 2003; Rubatto et al., 2006). The origin of this discrepancy
is unclear.
Another group of zircons, constituting roughly 50 % of the population show
“anomalous” patterns with enrichment in HREE, LREE, Th, Ti and P. These
analyses correspond to small domains observed in optical images with a yellow
colour and fine scale zoning (Fig. 2.12), resembling inherited cores. However
SHRIMP U-Pb analysis do not yield significantly older dates for such cores. Zir-
cons from sample B01-3 represent an interesting case where specific domains do
not have a distinct CL texture and age, but are evident optically and distinguish-
able by their trace element compositions. The LREE-enriched zircon patterns
can be explained by minute inclusions of an REE-rich mineral, for instance mon-
azite, though several observations cast doubt on this interpretation. Enrichment
in the LREE-HREE, Th, Ti and P occurs in these zones. The very high con-
centration of LREE and Th are characteristic of monazite, however Ti content
in monazite is very low (Chapter 1) and thus rutile inclusions are also possible.
101
Laser ablation time-resolved patterns indicate a homogeneous phase rather than
a two component mixture (e.g. zircon-monazite). The BSE images do not show
presence of inclusions in zircon cores. Therefore if inclusions exist they should
be smaller than 200 μm, in which case laser signal will be homogeneous. Zones
with anomalous composition in zircon from sample B01-3 can be interpreted as
inherited cores, in which Pb isotopic composition was reset during metamorphism
and it’s composition probably points to metamict state of the precursor zircon.
U-Pb SHRIMP dating of zircons from sample B01-3 returned several discor-
dant spots, high 204Pb and a relatively large spread of ages (MSWD = 4) and
outliers with significantly older or younger apparent ages. These features can
be attributed to a variable contribution of core domains with an enrichment in
REE, Ti and Th. Due to large spread of dates zircon dating results cannot be
interpreted as an age. Importantly zircons in sample B01-3 are completely meta-
morphic in age, and despite the relatively low peak metamorphic conditions of
this sample inherited ages are absent.
B94-333
B94− 333 Monazite:
Observations in thin section demonstrate that sample B94-333 monazite orig-
inates from the part of the sample with low-Ca garnet, and high-Ca garnet is
associated with allanite. Monazite homogeneous in BSE often has low HREE con-
tent. Monazite displaying zoning in BSE images sometimes shows higher HREE
contents. Several grains have oscillatory zoning, which is typical for melt-related
monazite, though their trace element composition is similar to other grains (Fig.
2.10). Monazite inclusion in garnet, analyzed in thin section has higher HREE
concentrations than matrix grain. Garnets in sample B94-333 have a continuous
decrease in HREE concentrations from core to rim. Variations in the HREE con-
tent of monazite may be attributed to the continuous growth of low-Ca garnet,
which preferentially extracted the HREE from the reactive bulk. High Sr con-
tents in monazite indicate that it crystallised at pressures exceeding the stability
field of feldspar (Finger & Krenn, 2007). During the decompression monazite
became unstable and was replaced by an assemblage of apatite, synchysite and
Th-bearing minerals (Section 3.6.1). Therefore monazite in B94-333 probably
was formed on prograde evolution, at peak conditions and at the beginning of
exhumation. Zoned and homogeneous monazite grains show overlapping range
of HREE concentrations, though zoned grains have a tendency to have higher
102
HREE content. Grains with different zoning and trace element composition show
no systematic age difference. Therefore it is concluded that different stages of
monazite growth occurred in the time interval of 528±7 Ma.
B94− 333 Zircon:
Variation in composition of garnet inclusions, CL structures, and trace ele-
ment composition of zircon indicates that zircons were formed from two distinct
bulk compositions in this banded sample. Ti-in-zircon temperatures are very sim-
ilar in both populations, thus indicating similar peak conditions. Zircons have
REE patterns with low HREE concentrations and small Eu anomalies. This is
evidence of their formation at high pressure conditions in the presence of garnet
and absence of feldspar. Several zircons contain inclusions similar to the man-
tles of low-Ca garnets, indicating their simultaneous formation. Zircon from this
sample was not dated.
B94-256
B94− 256 Monazite:
In sample B94-256 monazite has two growth zones: (a) homogeneous with
high HREE, and (b) low HREE with patchy and overgrowing homogeneous mon-
azite, both with prograde phengite inclusions. The decrease of HREE content
from the early to later generation can be explained by growth of monazite on the
prograde path in equilibrium with garnet, which progressively extracted HREE.
Then monazite survived peak metamorphic conditions, probably, because of ab-
sence of melting.
Several grains of allanite were found in the heavy mineral concentrate but
not observed in thin section. Monazite grains in thin section occasionally have a
corona of apatite, therefore suggesting allanite formation from monazite break-
down according to the reaction:
Monazite + fluid + Ca2+− > Allanite + Apatite.
It can be proposed that fluid which partly decomposed monazite also produced
high-Ca garnet present along rims and cracks (Fig. 2.4).
In this monazite, homogeneous domains yield an age of 515±7 Ma and patchy
zones date of 523±4 Ma, which are identical within the precision of measure-
ments.Therefore monazite dates prograde metamorphism of B94-256 at 521±6
Ma.
B94− 256 Zircon:
103
Zircons in B94-256 have two major growth zones: a core and a rim. From core
to rim, HREE and Ti contents increase. Inclusions in both domains of zircon cor-
respond to a high pressure association: Grt-Phe-Rt-Omph. The REE patterns
lack a negative Eu anomaly (Fig. 2.13) consistent with their HP origin. The
increase in Ti from cores to rims indicates increase of T and/or P during zircon
growth. The increase of REE content from core to rim may be related to a pres-
sure increase, but it is hard to explain by temperature increase because Zr/Grt
REE partitioning coefficients decrease with T (Rubatto & Hermann, 2007a).
One zircon core gave a 206Pb/207Pb age of 2867±72 Ma, demonstrating inher-
itance from a protolith. This grain also has a distinct trace element composition
with increasing REE abundance with atomic number and a high Th/U ratio of
0.14 (Fig. 2.13). All other zircons show very low Th/U (0.03–0.006), presumably
because Th was hosted by monazite. These low Th/U ratios of zircon from sam-
ple B94-256 differ markedly from samples which contained melt at some stage of
evolution. Zircon in samples B118A50 and B94-333 from this Chapter and from
B94-26 from Chapter 4 all have zones with high Th/U ratios, which indicate
monazite dissolution in the melt.
Zircon from B94-256 has distinct cores and rims, which do not have systematic
difference in age (cores: 522±6 Ma, and rims: 522±7 Ma). Therefore zircon date
is interpreted to pre-peak and peak conditions at 522± 6 Ma. Ages of monazite
and zircon are very similar.
B118A50
B118A50 Zircon:
In sample B118A50, zircons show concentric zoning and a trace element
composition similar to zircons from another UHP gneisse of the Barchi Kol’,
Kokchetav complex (see section 4.4.8). Zircon cores have low Th/U ratios, and
presumably grew on the prograde part of the PT path, when Th was hosted by
monazite. Mantles have highest Ti-in-zircon temperatures and high Th/U ratios,
indicating their growth at peak conditions, when monazite was dissolved in the
melt. The rims formed during exhumation. SHRIMP dating of zircons from
B118A50 resulted in large scatter of dates from 503±7 to 532±7 Ma and large
MSWD (Fig. 2.11). However this scatter is not related to the zircon’s texture
and/or composition.
104
Inheritance in monazite and zircon
Dating of monazite identified only metamorphic domains, and in zircon two cores
from two samples gave older, inherited ages. One zircon core from sample B94-
256 gave 206Pb/207Pb age of 2867 ± 72 Ma, though 206Pb/238U age is 1924 ± 23
Ma. A zircon core from sample B118A50 gave age of 617± 10 Ma.
Studied samples cover large range of conditions from medium grade sam-
ple B01-3 to UHP-UHT sample B118A50 and B94-256, also samples differ by
presence/absence of melt at some stage of rock evolution. Notwithstanding of
metamorphic conditions zircon and monazite in all investigated samples show
rare inheritance of older U-Pb ages. This is unusual and distinguishes Kokchetav
rocks from the common inheritance found in other UHP metamorphic rocks,
for instance in Western Gneiss Region, Norway (Corfu et al., 2003) and Dabie-
Sulu, China (Liu et al., 2001; Liou et al., 2002). This difference can be at-
tributed to metamorphic history of rocks. In Dabie-Sulu most common lithology
is orthogneisses, which are characterized by low water content and unreactive
assemblages. In Western Gneiss Region UHP metamorphism overprinted high
grade rocks of ancient basement, which were dehydrated by previous events. The
Kokchetav UHP rocks do not show any evidences of high grade metamorphism
prior UHP stage and thus protolith for UHP rocks likely were sediments of low or
no degree of metamorphism. Liberation of large amount of fluid during prograde
metamorphism and reaction of major minerals may be responsible for the rare
occurrence of inheritance in the Kokchetav UHP gneisses.
2.7.6 Origin of rocks with different PT path inside UHP
terrain
There are three potential scenarios that can explain how a rock without evidence
for UHP metamorphism can be inter-layered within the UHP unit:
A) The rock has been at UHP conditions but new minerals have not formed
at those conditions, and thus UHP stage was not recorded.
B) The rock has been at UHP conditions but it was completely recrystallized
and retrogressed and thus evidences for UHP stage were completely erased.
C) The rock never has reached UHP conditions, and was included into the
UHP block after UHP metamorphism.
Feasibility of scenario A is well demonstrated on HP rocks with incomplete
transformation of gabbros to eclogites in Western Gneiss Region and Zambia
105
(Engvik et al., 2001; John & Schenk, 2003). In such cases transformation of
gabbro to eclogite was controlled by the presence of hydrous fluid, and dry gab-
bros preserved their texture and composition, despite that their PT path was the
same as for eclogites. This scenario (A) is not applicable to the rocks from this
study because they all contain hydrous minerals and formed at high temperature
conditions. Whereas solid state recrystallization can be very lengthy process,
reactions in the presence of fluid/melt are fast, as evident from high-P high-T
experiments where equilibrium is attained after several days (Holloway & Wood,
1988).
Scenario B is more common than complete preservation of UHP associations:
in UHP gneisses and eclogites of the Kokchetav, Dabie-Sulu, and Western Gneiss
complexes typically inclusions in zircons are the only evidence left of UHP meta-
morphism (Sobolev et al., 1994; Carswell et al., 1999; Liu et al., 2002). Rocks
that mostly preserve UHP rock-forming minerals are very rare (the most no-
table examples are calc-silicate rocks from the Kokchetav and white schist from
the Kokchetav and Dora Maira). In such situations the obvious question is: how
does one distinguish between rocks, that never have been at UHP conditions from
rocks where mineral inclusions do not record UHP stage? In addition to zircon
inclusions, garnet and zircon trace element zoning provide complimentary infor-
mation to address this question. With increasing peak conditions garnet zoning
changes from intensive zoning both in major and trace elements, to homogeneous
major elements and zoning in HREE, and then to complete homogenization. Zon-
ing of garnet which experienced high temperature diffusional homogenization is
characteristically different from garnet which record prograde evolution and it
cannot be reproduced by any retrogression process. Zircon seems to record the
highest temperatures and thus provide basis for comparison of peak conditions of
various rocks. Therefore combined investigation of mineral inclusions with garnet
and zircon trace element zoning provides a much more reliable way to constrain
peak conditions than search for coesite inclusions. This approach demonstrates
that samples B01-3, B94-333 and B94-256 are not retrogressed UHP rocks, but
are rocks with different PT paths which reached lower metamorphic conditions
than the typical UHP gneiss B118A50.
Therefore I conclude that, for samples investigated in this study, the most
likely is scenario C of incorporation of the HP rocks into the UHP block on
exhumation.
106
2.7.7 Tectonic implications
In this study I obtained new U-Pb age data on the HP/UHP rocks from the
Barchi Kol’ UHP are of the Kokchetav complex. In this study monazite and
zircon from rocks with different PT paths were dated (Fig. 2.11, 2.18), whereas
previous geochronological studies were mostly concentrated on zircons from UHP
rocks of the Kokchetav complex (Claoue-Long et al., 1991; Hermann et al., 2001;
Katayama et al., 2001; Ragozin et al., 2009). In B01-3 dating of monazite con-
strains metamorphic peak at 528±8 Ma. According to monazite dating in B94-
333 metamorphism occurred at 528±7 Ma. Zircon ages in B118A50 vary from
503±7 Ma to 532±6 Ma and record prograde metamorphism, peak and exhuma-
tion. Ages of zircon and monazite from different samples coincide within the
precision of the measurements. It demonstrates, that regardless of peak condi-
tion achieved by the rock prograde, peak and exhumation metamorphism in the
Barchi Kol’ UHP area occurred over a relatively short time span, between 530
and 515 Ma. These estimates are consistent with previous estimates of the UHP
metamorphism and and exhumation on the Kokchetav UHP rocks obtained by U-
Pb dating at 528±5 Ma 2 (Hermann et al., 2001) or 519±8 Ma (Katayama et al.,
2001) and 524±25 Ma by Sm-Nd dating (Shatsky et al., 1999). Individual dates
in my study span from 503 to 538 and these scatter is substantially large than
analytical uncertainty. However because I do not find correspondence between
younger/older ages and earlier/latter growth zones on monazite and zircon I do
not assign any tectonic interpretation to these variations. Importantly, my results
constrain that exhumation occurred over a relatively short period of time, and
prograde metamorphism of the Kokchetav metasediments was contained within
the same time span.
In sample B94-256 monazite dates prograde metamorphism at 521±13 Ma
and zircon dates peak metamorphism at 522±6 Ma, thus B94-256 shows ages
within uncertainty estimates for other samples, but systematically younger than
other HP/UHP samples in the present study and dating of UHP gneisses by
Hermann et al. (2001). This age difference may suggests exhumation of sample
B94-256 after the main even of UHP exhumation in the Kokchetav.
Previous models propose that metamorphic belts with UHP rocks are com-
posed of terrains with a common metamorphic history (Kaneko et al., 2000; Liou
et al., 2002). Several facts indicate that typical Kokchetav UHP gneisses experi-
2Here error is propagated from ±2 Ma reported in Hermann et al. (2001) and 1% error onmeasurement (±5.3 Ma).
107
Figure 2.18: Comparison of U-Pb ages obtained in this study with results ofprevious investigations by Claoue-Long et al. (1991); Hermann et al. (2001);Katayama et al. (2001); Ragozin et al. (2009). D1–D4 ages of different genera-tions from study by Hermann et al. (2001). Black diamonds — maximum andminimum ages, blue diamonds — averages, with error bars for 95 % confidenceinterval or as reported in original studies.
108
enced similar peak conditions and PT paths. Firstly metamorphic diamonds are
very common in Kokchetav rocks and observed in several rock types (gneisses,
marbles, calc-silicate rocks). Thus, the majority of rocks achieved UHP condi-
tions. Secondly, diamonds from different rock types and samples have similar
range of nitrogen aggregation state (Cartigny et al., 2001; Nadolinny et al., 2006;
Sitnikova & Shatsky, 2009). These variations can be generally explained by vari-
ations in nitrogen content and aggregation kinetics (Nadolinny et al., 2006) and
different PT paths are unnecessary to invoke. Thirdly, the trace element compo-
sition of the majority of UHP gneisses is consistent with their partial melting and
melt extraction at similar PT conditions (Chapter 3). However, three samples
reported in this study have evidence of drastically different PT paths than the
typical UHP rocks. Sample B01-3 was collected within the UHP area and at
the distance of a kilometer from the main UHP locality. Samples B94-256 and
B94-333 were collected from the same drill cores as numerous UHP gneisses, at
the distance of tens or hundreds meters from other UHP samples. Therefore my
new data imply that, despite their close proximity, certain rocks in the UHP area
recorded different PT evolutions.
B01-3 was collected from the surface of the Barchi Kol’ area. The initial part
of the PT path of this sample (as recorded by garnet zonation), is characterised
by a very low temperature gradient, which only occurs in a subduction zone
environment. Then there is an almost isobaric increase of temperature, which
could occur if the rock was part of accretion prism which was heated and then
was incorporated into the UHP unit. This stage was followed by exhumation.
Source of the heat can be the warm part of the mantle wedge (discussed below)
or rising UHP rocks. Both isobaric heating and exhumation must have occurred
very quickly because garnet and phengite still preserve prograde compositional
zoning despite high temperature peak conditions.
B94-333 attained a lower peak temperature than a typical diamond-bearing
UHP gneiss such as B118A50. This sample it originates from the central part
of the Barchi Kol’ UHP unit and is closely spatially associated with diamond-
bearing UHP gniesses. The formation of this rock can be explained by the incor-
poration of lower grade rocks into UHP block during its exhumation. B94-333
shows an extreme degree of chemical heterogeneity: it contains layers with dif-
ferent assemblages and with garnets of varying compositions. Also elongation of
large garnets and pressure shadows indicate intense deformation. The heterogene-
ity of this sample can be explained if it was formed as tectonic mylonite and then
recrystallized. Heterogenious sediment protolith is less likely, because prograde
109
metamorphism generally results of homogenization of millimeter-centimeter scale
heterogeneities. Zircons from B94-333 form two populations with different zon-
ing, and probably originated from different parts of the sample. However Ti-
in-zircon are essentially identical in the two populations, thus indicating similar
peak conditions in the two parts of the sample. Sample B94-333 is likely to rep-
resent a rock derived from the wall of subduction/exhumation channel, because
it was trapped on the exhumation it reached substantially lower peak pressures
and temperatures than typical UHP gneisses. The intrusion of these wall rocks
happened by solid state flow, however, it can be facilitated by melting of the wall
rocks. Intrusion of colder rocks into the UHP slab can be additional mechanism
of cooling down UHP block on exhumation.
B94-256 shows peak temperatures similar to typical UHP gneisses, and it
probably reached similar peak conditions. However, this sample shows a con-
siderable difference from other UHP gneisses by the fact that the rock did not
melt and preserved some remnants of its prograde evolution. Melting might not
occur if two conditions were met simultaneously: (1) the rock was water under-
saturated, (2) the rock did not cross the upper limit of the phengite stability
field (Fig. 2.17). The low water content can by constrained by the presence of
Kfs inclusions in monazite, which formed on the prograde path. Several studies
concluded that Kokchetav rocks crossed phengite stability field (Hermann et al.,
2001; Auzanneau et al., 2006), however, B94-256 appears to be an exception.
The question is why? I propose that differences in exhumation paths of the rocks
are due to the temperature gradients in exhumation package during exhumation.
The majority of studies agree that the Kokchetav gneisses had a steep exhuma-
tion PT path with cooling on decompression (see Fig. 2.19 based on studies by
Dobretsov & Shatsky (2004), Zhang et al. (1997), Hermann et al. (2001), Auzan-
neau et al. (2006), Massonne (2003)). Interestingly this compilation demonstrates
that there is considerable variation in estimated temperatures of exhumation in
the range of pressures 30-10 kbar according different researchers. The cooling on
exhumation is attributed to the movement of UHP rocks in sudbuction channel
along a contact with colder rocks (Hermann et al., 2001).
The subduction zone is inevitably a region of large temperature gradients be-
cause the cold oceanic lithosphere constantly moves along the hot (above 1200oC)
mantle wedge (van Keken et al., 2002). The highest temperature gradients are
achieved at the boundary between mantle and top of the slab (van Keken et al.,
2002). Very fast exhumation of the Kokchetav UHP rocks favours the develop-
ment of large temperature gradients because dissipation of temperature gradient
110
Figure 2.19: Compilation of metamorphic paths of the Kokchetav UHP rocksproposed in different studies: 1– Dobretsov & Shatsky (2004), 2– Zhang et al.(1997), 3– Hermann et al. (2001), 4– Auzanneau et al. (2006), 5– Massonne(2003), 6– Katayama et al. (2001). Note that there is significant disagreementbetween different studies in the range of pressures from 30 to 10 kbar.
is a time dependent process.
2.7.8 Prograde PT path
There are several principal types of prograde PT evolution of metamorphic rocks
in a subduction zone. (1) Simultaneous increase of P and T where prograde PT
path is a straight line with a constant and small increase of T per increase of P.
(2) Upward bended PT path, where with increasing pressure influence of the hot
mantle starts to dominate over cold subducting front. (3) Downward bended PT
path (Cloos, 1982). Such a PT path can be similar to the exhumation PT path
of high pressure rocks (Peacock et al., 1994). (4) Kinked PT paths with generally
low T gradient but large increase of T over small increase of P at some stage of
prograde path.
The PT paths with kinks were predicted by several different models: (a)
111
flat subduction, where little increase of pressure with progressing warming of
the slab (Gutscher et al., 2000); (b) subduction with shear heating (Peacock
et al., 1994); (c) cold nose in mantle wedge. The low T-gradient part of PT
path occurs when slab goes along cold section of mantle wedge — “cold nose”
of mantle wedge (Honda et al., 2010; Gerya, 2011). When subduction brings
rock in contact with hot convecting mantle the temperature of slab increases
rapidly, particularly because of close contact of hot and cold rocks. Further the
temperature difference between slab and adjacent mantle decreases and gradient
of PT path flatten out. The “cold nose” of mantle wedge is gravitationally stable
due to low density of serpentinized peridotites.
Kinked PT paths were predicted by several independent models (Gerya, 2011;
Syracuse et al., 2010; van Keken et al., 2002), but apparently have not been ob-
served in nature, because of the difficulty of reconstructing prograde PT paths
in UHP rocks. Different models of kinked PT paths are based on different as-
sumptions and boundary conditions, but have a very important similarity: the
kink occurs at pressures within range 20-30 kbar, and temperature increases by
several hundred degrees.
Samples B01-3 and B94-256 both preserve prograde path, which in both cases
has stages with small temperature gradient and then large increase of tempera-
ture. This prograde path is equivalent to the kinked PT path. In B01-3 increase
of temperature occurred at 23-25 kbar and for B94-256 it is estimated at ≈30
kbar. These values are in agreement with predictions of geodynamic models.
2.7.9 Parallels with other UHP complexes
It is interesting to compare the new data about different PT paths in the Barchi
UHP terrain (Fig. 2.17) with the available data about variability of PT paths in
other major UHP terrains (Fig. 2.1). In the Western Alps UHP complexes of
the Dora Maira and Lago di Cignana, the UHP rocks are confined to relatively
small slab blocks with thicknesses of hundreds of m metres and width/length
of kilometers (Chopin, 1987; Reinecke, 1998; Beltrando et al., 2010). In the
Lago di Cignana complex UHP blocks have complex contacts with country rocks
(Forster et al., 2004) and form of the UHP blocks can be considered as boudens
within lower pressure country rocks. These structures probably can be considered
as analogous to the Kokchetav and they are more apparent, because of better
exposure in the Alps.
In the Western Gneiss Region (WGR), Norway, UHP rocks are observed
112
in several large areas. The questions of whether the boundary between UHP and
HP rocks is a tectonic displacement or simply a gradual metamorphic transition
and what is the scale of UHP metamorphic terrains have long been a matter
of debate. Smith (1980) proposed that the whole region is a tectonic melange
and the HP-UHP rocks are foreign to the host rocks. Recent models favor a
structure of the region without major tectonic disturbances with a consistent
increase in metamorphic grade from east to west culminating in the UHP localities
(Carswell et al., 1999; Cuthbert et al., 2000; Root et al., 2005; Hacker et al., 2010).
Major element zoning of garnet is used in the WGR as an important criterion to
distinguish HP and UHP rocks, with HP rocks often preserving prograde zoning
while UHP do not (Cuthbert et al., 2000). Results from this study demonstrate
that trace element composition of zircon also can be useful in discrimination of
HP/UHP rocks.
In China, the large Dabie-Sulu UHP belt contains UHP terrains with differ-
ent peak conditions and there are trustworthy reports of non-UHP rocks inside
UHP terrains. Wang & Cong (1999) concluded that on the scale of tens of
meters peak pressures and PT paths vary in eclogites of the Dabie Mountains.
HP/UHP associations was divide into two types: (1) Eclogites in a matrix of
metasedimentary rocks, where both eclogites and metasediments were metamor-
phosed at UHP conditions. (2) Eclogites in matrix of granitic orthogneisses,
where orthogneisses were metamorphosed only at amphibolite facies conditions
and eclogites at UHP. However, based on findings of coesite in othogneisses Liou
et al. (2002) concluded that all lithologies within the Dabie UHP terrain experi-
enced UHP metamorphism and have identical PT paths. Important contributions
to this discussion are series of systematic studies of mineral inclusions in zircons
from drillholes samples of the Chinese Continental Scientific Drilling Project (Liu
et al., 2001, 2002, 2007). In the first pilot drilling CCSD-PP1 it was observed
that a 40 meters thick layer of garnet and epidote bearing granite gneisses from
which three samples did not contain any coesite inclusions in zircon (Liu et al.,
2001). In another pilot drilling CCSD-PP2 samples of paragneiss without coesite
from the depth of 425 meters bounded above and below by UHP rocks (Liu et al.,
2002) were reported. On the other hand, the study of the 5.5 km long main drill-
hole CCSD-MH of the Chinese Continental Scientific Drilling Project found that
all 137 samples contained inclusions of coesite (Liu et al., 2007). Unfortunately
no additional information (data on mineral inclusions, geochronology, mineral
zoning, etc.) is reported about the non-coesite samples. The proposed variant
of origin of the non-UHP gneisses is granitic intrusion at low depth after UHP
113
metamorphism (Liu et al., 2001).
Rocks with completely different evolution were found by Zhou et al. (2008),
who reported finding greenschist facies quartzites within the Sulu UHP metamor-
phic belt. A U-Pb study of detrital zircons from these quartzites revealed that
they represent sediments derived from the North China Craton (NCC) whereas
all of UHP rocks originate from the South China Craton (SCC). Ar-Ar dating of
mica provided an age of 266±1 Ma, which is older than the main stage of UHP
metamorphism at 210-255 Ma in the Sulu belt (Liu et al., 2001). Ar-Ar ages are
interpreted to date regional metamorphism, and appear to be undisturbed by
emplacement into the UHP terrain. The origin of the greenschist facies quarzites
was explained by introduction of slices of low grade NCC rocks between UHP
SCC blocks during exhumation.
These studies demonstrated that non-UHP rocks form small but important
fractions in UHP sequences and confirm the possibility of solid state intrusion of
rocks with lower grade into UHP blocks during exhumation as proposed for sam-
ples B01-3 and B94-333. The Dabie-Sulu UHP complex is substantially larger
than the Kokchetav complex: the Dabie-Sulu belt is >1000 km in length, several
tens of kilometers wide and at least 4.5 km deep (Liu et al., 2007; Zhang et al.,
2009). The subduction-exhumation of Dabie-Sulu took longer than that of the
Kokchetav (210-245 Ma Zhang et al., 2009). The large size of the UHP block
in Dabie-Sulu UHP belt should favour development of temperature gradients in
the UHP complex, whereas long duration of UHP metamorphism and exhuma-
tion assisted dissipation of temperature gradients. Therefore it is difficult to
predict whether temperature gradients were present in the Dabie-Sulu complex
during UHP metamorphism. These issues should be investigated by modeling
and geothermobarometry.
Diamond-bearing quartzofeldspathic rocks are present in Erzgebirge, Bo-
hemian massif, Germany (Massonne, 1999). These rocks have a pelitic pro-
tolith, form relatively small lenses (100X10 m in size) and are hosted by or-
thogneisses, which themselves lack evidence of UHP metamorphism. Massonne
(2003) proposed that the quartzofeldspathic rocks of Erzgebirge reached condi-
tions of 1200oC and 60 kbar. According to this model the rocks were essentially
magmatic rocks: a mush of crystals and melt when exhumation occurred. How-
ever, there are no observations of intrusive contacts of the UHP rocks with their
host orthogneisses or xenoliths, and as such their relationship with the host rocks
remain uncertain. The Erzgebirge has a structure with relatively thin lenses of
UHP rocks in the non-UHP matrix, which can be considered as similar to the
114
association of rocks with different PT paths in the Kokchetav. Substantial dif-
ference is that in the Erzgebirge the fraction of the UHP rocks relative to HP
matrix is lower than in the Kokchetav.
2.8 Conclusions
Detailed petrographic and geochronological investigation of metasedimentary rocks
from the Barchi Kol’ UHP area in the Kokchetav complex demonstrate examples
of rocks with PT paths within UHP terrain:
# Sample B01-3 has not been molten and achieved relatively low peak con-
ditions of 710oC and 24 kbar. The sample experienced PT path with initial
increase of pressure and then increase of temperature, which can be considered
as evidence of kinked PT path during subduction. This sample likely derived
from another part of the subduction channel and was embedded into the UHP
block on exhumation.
# Sample B94-333 rock represents former mylonite of rocks of different com-
position which experienced peak metamorphism at 800–900oC and 20–30 kbar in
subduction zone and then was trapped by UHP rocks during exhumation.
# In B94-256 peak conditions were close to the peak of UHP metamorphism,
approximately 1000oC and 45 kbar. However, on exhumation the decrease of
temperature was faster than in the majority of UHP gneisses and the rock avoided
partial melting. Different rate of cooling of UHP rocks during exhumation can be
the result of temperature gradients in the subduction zone/exhumation channel.
Timing of the metamorphism in the investigated samples (521–528 Ma) was
similar to the metamorphism of UHP rocks, demonstrating that deep subduction,
peak metamorphism and exhumation occurred over a relatively short period of
time. Two samples demonstrate evidences for kinked prograde PT paths where
initially pressure increased and then temperature rose rapidly. Such PT evolution
of suducting rocks is in a good agreement with modern models of a subduction
zone PT geotherms.
This study demonstrates the importance of accessory minerals not only for
geochronology but also for reconstruction of the PT path of the HP/UHP rocks.
In major minerals, information about prograge evolution and peak conditions
is limited to trace element zoning in garnet. In accessory minerals, trace ele-
ment compositions and inclusions provide much more detailed record and may
be the only evidence of prograge evolution and peak conditions. This study also
documents the application of garnet trace element zoning for discrimination of
115
HP/UHP rocks with different histories. Combination of garnet trace element
zoning with trace element geochemistry of resistant accessory minerals produces
a detailed picture of subduction metamorphism.
The main finding of this study is that UHP complexes have a more compli-
cated internal structure than it is usually thought. Within the UHP block of the
Kokchetav complex rocks with completely different metamorphic histories are
associated on the 10 meter scale. Together with works Rubatto et al. (2011) that
proposed yo-yo subduction with repeated cycles of subduction and exhumation
within the same rock unit, these finding demonstrate the complexity of tectonic
in subduction zones.
2.9 Supplementary materials
Tables with major and trace element compositions of minerals, and results of
U-Pg dating are resented in electronic supplementary material on a CD.
116
Table 2.1: Comparison of various petrological parameters of studied samples.
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Chapter 3
The geochemistry of Ultra High
Pressure anatexis
3.1 Introduction
At active plate margins, oceanic lithosphere and its sedimentary cover are sub-
ducted into the mantle and undergo heating, dehydration, partial melting and
mixing with mantle material. Metasediments are the most fertile component of
the subducting lithosphere: they have the lowest melting temperature, and are
thus able to produce high melt fraction during subduction. This liberates high
abundances of trace elements which are characteristic of subduction zone mag-
matism (Plank & Langmuir, 1998; Plank, 2005). Subduction zone processes are
studied by means of experimental and numerical modeling, by geophysical meth-
ods and through geochemical investigations of arc magmas. Restites after partial
melting at great depth descend further into the mantle potentially forming spe-
cific mantle reservoirs (Zindler & Hart, 1986; Chauvel et al., 2008), and can be
a complementary source of information on subduction zone processes. Unfortu-
nately, such restites of subducted metasediments are rarely available for study
because they return to the surface only at special circumstances.
Ultrahigh Pressure metamorphic (UHP) rocks have undergone subduction to
great depths (where coesite is stable) and were subsequently exhumed to the
surface. These rocks can be used to study elemental fractionation in subduction
zones. HP/UHP rocks retain their trace elements after fluid loss at P-T conditions
up to 40 kbar, 700oC (Hermann, 2002a; Spandler et al., 2003). Trace elements
were mobilised during exhumation of UHP rocks in Dabie-Sulu complex (Zheng
et al., 2011b; Hermann & Rubatto, 2012). Higher temperatures are necessary for
efficient extraction of incompatible elements by melting in subduction zone.
The Kokchetav complex in Kazakhstan is ideal for the study of melting at
UHP conditions, as these rocks experienced UHP metamorphism at peak con-
117
118
ditions of 45–60 kbar and 950–1000oC and contain abundant meta-sedimentary
rocks that have undergone partial melting (Sobolev & Shatsky, 1990; Shatsky
et al., 1999; Hermann et al., 2006). Partial melting of the Kokchetav rocks was
proposed on the basis of bulk rock geochemistry (Shatsky et al., 1999). The
Kokchetav UHP gneisses are depleted in K, Na, Si and LREE (Shatsky et al.,
1999, 1995, 2006b). This Chapter presents new data on bulk rock major and
trace element compositions of the metasedimentary rocks of the Kokchetav com-
plex. In order to quantify the mobility of major and trace elements during UHP
metamorphism the compositions of UHP gneisses is compared with lower grade
samples, considered as potential protolith for UHP gneisses. The composition of
the UHP gneisses is interpreted in the light of recent experimental studies (Her-
mann & Rubatto, 2009, Chapter 1) providing first constraints on the melting
history, mineral assemblages during melting and melt extraction parameters of
the Kokchetav metasediments.
3.2 Geological background
The Kokchetav metamorphic belt is located in the central part of Eurasia (Fig.
3.1) within the large Paleozoic Central Asia Fold Belt. Magmatic rocks are
widespread in the Kokchetav metamorphic belt (Shatsky et al., 1999) and are
dominated by granites and granitic orthogneisses (gneisses of granitic composi-
tion). They include granites, that predate and postdate UHP metamorphism.
Gneisses with broadly granitic composition (e.g. composed essentially of quartz,
feldspars and with low content of mafic minerals) are abundant in the Kokchetav
complex and they are considered to represent the basement of the Kokchetav
continent (Dobretsov et al., 1995). The important localities include outcrops
near Kulet lake (Kushev & Vinogradov, 1978), shores of Chaglinka water reser-
voir (Turkina et al., 2011) and outcrops along the Kumdy Kol’ lake (Kushev &
Vinogradov, 1978). The U–Pb dating of zircons from gneisses provide ages 1128
± 12 and 1169 ± 12 Ma (Letnikov et al., 2007; Turkina et al., 2011). Turkina
et al. (2011) demonstrated that the granites are of A-type and attributed them
to crustal melts from continental rift zone.
Among the post UHP granites is the large Zerenda batholith (460–440 Ma)
which composes the southern part of the Kokchetav complex (Shatagin et al.,
1999, 2001). Several smaller intrusions are widespread in the Kokchetav complex
and some of these granites are characterized by high fluorine contents, the pres-
ence of topaz and high contents of trace metals (Letnikov, 2008). Shatsky et al.
119
(1999) proposed that trondhjemite veins within Kumdy Kol’ UHP terrain were
formed by melting of eclogites, however other magmatic rocks with ages close to
the time of UHP metamorphism are absent absent outside of the UHP terrains.
Partial melting of some of the UHP metamorphic rocks of the Kokchetav
complex was proposed by Shatsky et al. (1995, 1999). These authors concluded
that Kokchetav rocks melted on the basis of geochemistry, and in particular from
depletion in LREE. Also, experimental studies indicate that at peak conditions
Kokchetav was clearly above the wet solidus and close to the fluid-absent melting
curve (Hermann & Green, 2001; Auzanneau et al., 2006, see Chapter 2 and Fig.
2.17 there). Finally unequivocal evidence of melting were found as glass/fluid
inclusions in diamonds (Hwang et al., 2006) and in polyphase inclusions in UHP
minerals (Korsakov & Hermann, 2006).
3.3 Analytical methods
For bulk rock analysis and mineral separation the samples for this study were
ground in either a tungsten carbide mill or in an alumina mill to a grain size of
< 400μm. Approximately 75 % of the sample was used for mineral separation
by the conventional method of heavy liquids and magnetic separation and the
remained sample was powdered to < 25μm in an alumina oxide mill and used
for bulk-rock analysis. Major/manor elements (Na, Mg, Al, Si, P, S, K, Ca,
Ti, Mn, Fe, S, Cl, F) concentrations were determined by XRF of fused discs
with a Phillips (PANalytical) PW2400 X-ray fluorescence spectrometer in RSES
at ANU by Ulrike Troitzsch. The lithium borate fusion discs were prepared by
fusion of 0.27g of dried sample powder and 1.72g of “12–22” eutectic lithium
metaborate-lithium tetraborate. The major elements were calibrated against 28
international standard rock powders. Another set of samples was measured by
Axios from PANalytical with a 2.4kWatt Rh X-ray Tube and by in the Central
Analytical Facility of Stellenbosch University, South Africa operated by Esme
Spicer. 1 g of sample powder was mixed with 10 g of Claisse flux and fused in
M4 Claisse fluxer for 23 minutes.
Trace elements were analysed by LA-ICP-MS of fused discs at RSES, ANU,
using a pulsed 193 nm ArF Excimer laser with 100 mJ energy at a repetition
rate of 5 Hz (Eggins et al., 1998) coupled to an Agilent 7500 quadrupole ICP-
MS. Laser sampling was performed in an He–Ar–H2 atmosphere using spot size
between 105–137 μm. Data acquisition was performed by peak hopping in pulse
counting mode, acquiring individual intensity data for each element during each
120
Figure 3.1: Simplified geological outlines of the Kokchetav complex from Do-bretsov et al. (1995, 1998); Korsakov et al. (2006). Location of the Kokchetavcomplex in Northern Kazakhstan is shown in the inset. Simplified setting of UHProcks in Kumdy Kol’ Western and diamond-bearing (Dia) and Kulet, eastern andcoesite-bearing (Coe) sub-domain in the Kokchetav Megamelange Domain (af-ter Dobretsov et al. (1998). Red dots show locations of sampling of non-UHPsamples.
mass spectrometer sweep. A total of 60 seconds, comprising a gas background of
20–25 seconds, were performed for each analysis. The LA data were processed by
an Excel spreadsheet created by Charlotte Allen. Compositions were calculated
using NIST 612 as external standard and SiO2 as the internal standard. The
resulting data are an average of 4–5 ablation spots; 1 standard deviation from
the average is < 5% relative for most elements. BCR glass was measured as
secondary standard and deviation from values by Norman et al. (1998) was < 5%.
In this study I investigate relatively small samples (30–50 g) and the main
focus is on the behaviour of trace elements, which are mainly stored in accessory
phases. Given the small grain size of these accessory minerals, representative
samples must be ensured. In geochemistry this is especially relevant for elements
like noble metals, which are concentrated in nuggets in ore samples. According
to Potts (1996), in samples with a grain size of 30μm and containing 50 ppm
Zr, 0.1 g of sample will be sufficient to achieve 5 % σ sampling precision. In the
present chapter compositions of UHP gneisses and protolith metasediments are
121
compared. Since variability of composition of these rocks is much larger than
5-10 % I conclude that relatively small sample size should not be a problem for
trace elements in these samples.
3.4 Sample description
This study aims to determine the effect of UHP melting on the major and trace
element composition of metasedimentary rocks. Therefore bulk compositions of
two types of samples were investigated: UHP gneisses and non-UHP metasedi-
ments and orthogneisses, which are considered as potential protoliths for UHP
gneisses (Fig. 3.2).
3.4.1 Non-UHP samples
B01− 3: garnet-micaschist, collected as loose stone from the surface in the Barchi
Kol’ area (for detailed description, petrology and geochronology see Chapter 2,
and map 2.2). The sample consists of: Grt, Qtz, Ky, Phe with a small amount
of retrograde biotite and accessory minerals: Rt, Mnz, Ap and Zrn. The sample
has a weak foliation defined by Phe.
B94− 256: garnet-micaschist from drill hole in the Barchi Kol’ area (Fig. 2.2,
for the detailed description, petrology and geochronology see Chapter 2, and map
3.2). The rock has a homogeneous texture with a weak foliation and is composed
of major minerals Grt, Qtz, Phe, Bt, Kfs and accessory minerals, which include
Rt, Zrn, Ap, Aln and Mnz. Garnet grains are ≈2 mm in size. Biotite forms large
grains in the matrix and discontinuous rims around garnet.
B94− 333: garnet-gneiss from drill hole in the Barchi Kol’ area (detailed
description, petrology and geochronology see Chapter 2, and map 2.2). The
rock is composed of Grt, Bt, Phe, Ky, Qtz and accessory Rt, Zrn, Ap, Mnz
and sulfides. The gneiss texture is marked by thin leucocratic layers enriched
in quartz and feldspars. One layer contains a few large grains of garnet (up to
9× 5 mm). Large garnets contain inclusions of rutile and zircon in the core and
mantle and monazite inclusions close to the rims.
B95− 99: sample from drill hole 121 in the Barchi Kol’ area. It is a mica-
schist cut by a leucogranite vein with large phengite grains. The matrix consists
of Grt, Phe, Bt, Rt, Kfs, Pl and accessory Ap, Mnz, Aln. A granite vein is 1–2
cm thick and is composed of large phengite grains (5 mm long), Ab, Kfs, Qtz
and Ap.
122
Figure 3.2: Map of Bachi Kol’ area (from (Korsakov et al., 2002)) and UHP sam-ples localities. Symbols: 1 – Quartzites, sericite-quartz schists (Kokchetav For-mation); 2 – Sericite-chloritic, carbon-carbonate schists, limestones (Sharyk For-mation); 3 – Kyanite, clinopyroxene, clinozoisite, mica-bearing gneisses; garnet-pyroxene and silicate-carbonate rocks; amphibolites (Tectonic Unit 1); 4 – Py-roxenites, micaceous pyroxenites, carbonatites (Barchi Massif, Krasnomai alkali-utrabasic complex); 5 – (1) Leucocratic fine-grained granites, alaskitic granites;(2) orthogneisses; 6 – Eclogites; 7 – Geological boundaries; 8 – Tectonic distur-bances.
123
Samples K03-1 ans K03-2 were taken from an outcrop on the shore of lake
Kumdy Kol’ to the east of the entrance of Kumdy Kol’ exploration gallery, from
a sequence of gneisses and micaschists with lenses of eclogites.
K03− 1: orthogneiss composed of major minerals Pl, Kfs, Qtz and accessory
Aln and Zrn.
K03− 2: micaschist with a strong foliation. The sample contains Qtz, Phe,
Ky and accessory Zrn and Mnz. Phengite grains are up to 3 mm in length, they
are bent and deformed.
CHAG: the sample originates from the outcrop near Chaglinka river. It is a
micaschist with a porphyroblastic texture and big (up to 3 cm) rare garnet grains
surrounded by biotite rims. The matrix is foliated and composed of Phe, Ky, Bt
and Qtz.
Q03− 11: “metaconglomerate” gneiss from the outcrop in Enbek-Berlyk area.
The sample has a coronitic texture with large layered ovoids and a weak foliation.
The rock is composed of Grt, Ky, Bt and accessory Rt, Ap, Mnz. In the centers
of ovoids there are grains of corundum, then Ky + Rt, then a layer of Grt + Qtz
with Ab in-between ovoids.
Q03− 12: “metaconglomerate” gneiss from the outcrop in Enbek-Berlyk area
(Fig. 3.2). The rock is composed of Ky, Grt, Bi, Qtz and accessory Rt, Ap, Ilm,
Mnz.
Ky98− 3: cordierite fels from the Daulet suite outcrop in Kulet area (Fig.
3.2). The rock is black and has a massive texture. The rock contains mainly Crd
and Qtz with minor Bt and Kfs. Accessory phases are Ilm, Sp and Mnz.
Ky98− 18: orthogneiss from an exposure in the Kulet area (Fig. 3.2). This
is the host rock for the Kulet white schists. The sample contains: Qtz, Kfs, Pl,
Bt and Zrn. The sample shows a folded texture with a characteristic wavelength
of about 5 cm.
G03: mica-schist from Sulu-Tjube. The sample is composed of Grt, Phe,
Bt, Qtz, Ky and accessory Rt, Zr, Tur and Mnz. The sample has a strong
lineation formed by large wavy flakes of phengite. Garnet grains are rounded,
approximately 1 mm in size, with abundant mineral inclusions. Sample B98-3
represents low pressure contact metamorphism.
The samples of non-UHP rocks originate from various localities of the Kokchetav
complex and were selected in order to represent the spectrum of metasedimentary
and felsic rocks in the Kokchetav metamorphic belt 3.2. The rocks cover a wide
range of metamorphic conditions. Samples Ky98-18 and K03-1 are orthogneisses,
which are metamorphosed magmatic rocks of granitic composition. Samples K03-
124
2, CHAG, Q03-11, Q03-12 and G03 were metamorphosed at eclogite facies. Four
samples (B01-3, B94-333, B94-256, B95-99) originate from the Barchi Kol’ UHP
area. Though these samples originate from the same location as UHP gneisses
they had different PT paths than UHP gneisses (Chapter 2). Samples B01-3 an
B94-256 have not been melted. Samples B94-333 and B95-99 show some textu-
ral evidences of melting (gneissose texture and vein of Qtz-Fsp-Phe leucogranite
in sample B95-99) but they have achieved lower peak conditions than the UHP
gneisses (Chapter 2) and their composition is similar to other non-UHP metased-
iments. Therefore these rocks are also suitable to constrain the protolith to UHP
gneisses.
3.4.2 UHP samples
The UHP garnet-biotite gneisses were selected with the intention to represent
typical UHP gneisses from the large collection of drill cores from the Barchi Kol’
area. Most of the samples are fresh, as is evident from the presence of non oxidized
sulfides. Evidences of UHP metamorphism of these rocks include presence of
diamond inclusions in some zircons, (e.g. B11A50), high garnet contents, absence
of garnet zoning or presence of only retrograde zoning and their origin from the
UHP terrain.
B94− 57 garnet-gneiss from a drill hole in Barchi Kol’ area. It contains Grt,
Kfs, Pl, Qtz, Bt and accessory Rt, Zrn, Ap and sulfides. Spindle shaped garnet
grains are elongated in the direction of the lineation (Fig. 3.3). The garnets lack
inclusions and are partly replaced by chlorite.
B94− 29 garnet-biotite gneiss from drill hole in Barchi Kol’ area. The sample
contains Grt, Bt, Ky, Qtz, Kfs and accessory Zr and Ap. The sample shows a
significant foliation and folding with melanocratic layers composed of Grt and Bt
and light layers containing Grt with feldspars and Qtz (Fig. 3.3).
B94− 355 garnet-gneiss from the drill hole in Barchi Kol’ area. It contains
Grt, Ky, Kfs, Qtz, Bt. Garnets have a light pink colour, irregular shape and
contain polyphase inclusions together with rounded inclusions of Ky and Rt.
The matrix is composed of feldspars and Qtz. Kyanite grains are rare and have
a corroded shape. Biotite is also rare and represents a retrograde mineral near
garnet.
B94− 60: garnet-biotite gneiss from drill hole in Barchi Kol’ area. The sam-
ple contains Grt, Bt, Qtz, Kfs, Pl and accessory Zr and Aln. The sample shows
a significant foliation (Fig. 3.3).
125
Samples B118A1−B118A17 were collected from the depth of 58.2–61 m of
the drill hole in the Barchi Kol’ unit. These 8 samples represent the meter-scale
heterogeneity of the UHP gneisses. Gneisses have homogeneous texture some-
times with a weak lineation, formed by aggregates of leucocratic and melano-
cratic minerals. The samples are composed of fine grained Grt, Kfs, Pl, Bt, Phe
and Qtz. Accessory minerals are Zrn, Rt, Ap and sulfides. Garnet crystals are
small (1–2 mm), cracked and commonly have biotite rims.
3.5 Results
3.5.1 Classification of non-UHP Kokchetav rocks
I selected 12 samples as representative for the Kokchetav non-UHP felsic-metapelitic
rocks. The most apparent features of the major element compositions of these
samples is their high K2O content and very high K2O/Na2O ratios, ranging from
2 to 32. Based on the major elements compositions several groups can be defined:
Orthogneisses. Samples Ky98-18 and KK03-1 are composed mostly of quartz
and feldspars. Sample KK03-2 has a high content of phengite but its composition
is virtually identical to samples Ky98-18 and KK03-1 and it probably represents
the product of hydration of orthogneisses. These samples have a low content of
“mafic” components (CaO+FeO+MgO 2.6–3.2 wt.%), high Na2O + K2O (4.7-7
%) and SiO2.
Metasediments. 7 samples have a high content of “mafic” components (CaO
+ FeO + MgO 5-17 %) together with a high “felsic” component Na2O+K2O (3.3–
4.8%) and high Al content. These features are typical for metasediments (Taylor
& McLennan, 1985; Behn et al., 2011; Hacker et al., 2011), and they cannot
have a magmatic protolith because typical granites have much lower content
of “mafic” components, whereas high K2O mafic magmatic rocks are rare and
always alkaline.
High− Al gneisses. Samples Q03-11 and Q03-12 from Enbek-Berlyk have
conspicuous major element compositions with very high Al2O3 (≈ 24wt.%) and
low SiO2 (54− 59wt.%).
Based on the Na content, metapelites can be subdivided into two additional
groups:
Low −Na metasediments. Samples B94-256, B01-3, Ky98-3 and G03 have
CaO <1 wt.%, Na2O < 0.5 wt.% and K2O/Na2O > 10. They contain 3.1–4.6
wt.% of K2O, 63–76 % SiO2 and high Fe2O3total 6–8.4 wt.%. They have low P
126
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Figure 3.3: Scanned images of the UHP gneisses. The gneisses are rich in garnetand show gneissic texture: dark areas are mostly composed of Grt+Bt and lightcolored bands are enriched in Fsp+Qtz.
127
(0.02–0.08 wt.%), Mn (MnO < 0.25 wt.%) and Sr (6–25 ppm) content. These
samples are also characterised by moderate contents of Rb (102–230 ppm) and
Ba (240–620 ppm).
High−Na metapelites. Samples CHAG, B94-333 and B95-99 contain Na2O >
0.4 wt.% and 2 < K2O/Na2O < 4. They typically contain CaO >1 wt.% though
sample CHAG has only 0.4wt.%. Sr content is higher than in the previous group
at 40–51 ppm.
All non-UHP metapelites samples have a concave upward REE pattern with
a small negative Eu anomaly. According to the trace element compositions of
the samples, they can be classified into two groups, which also are distinguished
by their Ta and Nb contents:
Metasediments with high− LREE content of 250–280 ppm∑
LREE. These
samples have high Th (25–30 ppm) and high Th/U ratios 5–8 (in comparison with
average upper crustal Th/U=4 (Taylor & McLennan, 1985)). These samples also
havean elevated content of Nb (20–25 ppm) and Ta (2.1–2.8 ppm). Both samples
of high-Al gneisses correspond to this category, but there is no correspondence
between Na and LREE concentrations.
Metasediments with moderate− LREE content of 110-140 ppm∑
LREE.
Concentrations of Nb (10-17 ppm) and Ta (1.1–1.4 ppm) are lower than in the
previous group.
3.5.2 REE minerals in UHP gneisses
Accessory minerals rich in LREE, Th and U were identified using SEM investi-
gation (BSE images). These minerals are relatively easy to find, because they
contain the heaviest elements of the periodic table and hence appear very bright
on BSE and SEM images.
In the majority of non-UHP samples, the dominant LREE phase is monazite
(samples B01-3, G03, Ky98-3, Q03-11). Usually monazite forms well developed
large grains 100–200μm in size. In sample K03-2 monazite forms fine grained
opaque aggregates. Allanite is the main host for LREE in only sample K03-1.
In sample CHAG kyanite and garnet contain polyphase inclusions of monazite-
apatite-bearthite Ca2Al(PO4)2(OH). Monazites contain high Sr (up to 2 wt.%
SrO). Similarly in sample Q03-12 monazite is present either as separate grains
or in aggregates of apatite and monazite. It is likely that monazite in these as-
sociations was formed from bearthite (Chopin et al., 1993; Scherrer et al., 2001).
In the sample B94-256 most LREE is hosted by monazite, but allanite is also
128
present (see section 2.7.5). A rich association of REE minerals was observed in
the non-UHP sample B94-333 from Barchi Kol’. It contains monazite inclusions
in garnet and matrix grains. In the sample there are several reactions involving
the LREE-bearing minerals. One monazite grain in the matrix is particularly
important. It is located between an apatite crystal and garnet. Apatite con-
tains numerous inclusions of Th minerals, probably (ThSiO4) (Fig. 3.4f). In
the matrix there are grains of allanite and a relatively rare mineral synchysyte
(La, Ce)Ca(CO3)2F with low Th content (Fig. 3.4e).
In the UHP gneiss LREE, Th and U minerals are rare. Sample B118A50
contains an aggregate of phengite with 3.43 Si pfu, thorite/huttonite (ThSiO4)
and some phosphates with variable composition (Fig. 3.4). Due to its morphol-
ogy, this aggregate is interpreted as an allanite pseudomorph (Fig. 3.4a-b). This
rock also contains large rutile grains with inclusions of uraninite (Fig. 3.4d). Ru-
tile is surrounded by a symplectite of Rt-Phe with Nd-rich monazite (Fig. 3.4c).
In the UHP gneiss B94-60 allanite grains are present in the matrix. A big titan-
ite grain has been partly replaced by a symplectite of rutile with phengite and
contains small monazite grains (Fig. 3.4h). In other UHP samples I did not find
LREE minerals, primarily due to low concentration of LREE in these depleted
gneisses, but also due to presence of sulphides, which complicate the search for
bright grains on SEM images.
3.5.3 Compositions of UHP gneisses and comparison with
potential protolithes
Major elements
UHP gneisses have a restricted range of compositions, with 60–66 wt.% SiO2
(sample B94-29 has 71 wt.%), and 10–15 wt.% Al2O3 (Fig. 3.5, Table 3.1).
The UHP gneisses exhibit high K/Na ratios which are also characteristic of the
non-UHP Kokchetav rocks (Table 3.2). The UHP gneisses contain higher Na2O
(0.4–1.2 wt.%) than the majority of non-UHP metasediments with <0.4 wt.%
of Na2O and only the non-UHP samples from the Barchi Kol’ (B94-256, B94-
333 and B95-99) have Na2O content comparable with UHP gneisses. The UHP
gneisses contain 2.0–5.1 CaO wt.%, which is comparable to only two non-UHP
metasediments (B94-333 and B95-99). These samples also have elevated Na2O
content.
129
Figure 3.4: Photograph in transmitted light (a), and BSE images (b–h) of REE,Th and U minerals in the studied samples. (a–d) sample B118A50. (e–f) sampleB94-333. (g–h) sample B94-60. Mnz – monazite, Aln–allanite, Sy – synchysyte(La, Ce)Ca(CO3)2F
130
Figure 3.5: Major element compositions of potential protolith rocks (metasedi-ments and orthogneisses), and UHP gneisses from this study (TS) and Shatskyet al. (1999) (S99) . Subtraction of melts from HP experiments by Hermann &Spandler (2008) (HS) from the estimated protolith composition (KMC) producesthe trends shown by the line, which are in good agreement with observed compo-sitions of UHP gneisses. 1000–45 denotes 1000oC, 45 kbar, etc. — conditions ofthe experiments which produced that melt. Crosses on the melt extraction trendline denote increment by 10 % of mass extracted by melt from the protolith.
131
Figure 3.6: Comparison of elements abundances between the UHP gneisses andnon-UHP samples normalized to chondrite concentrations: (a) LREE, Th and U,(b) LILE, (c) HFSE, (d) transitional elements.
REE, Th and U
LREE, Th and U are strongly depleted in the UHP gneisses relative to potential
protolith samples and with respect to average composition of the continental
132
crust (Fig. 3.6). REE pattens of UHP gneisses are highly variable, ranging from
those which are concave upward to patterns with depletion in Pr and Nd relative
to La and Sm, to patterns which have flat and very low LREE. All pattens have
a similar Eu anomaly Eu/Eu∗=0.50–0.77, which is similar to the negative Eu
anomaly of protolith metasediments and to the typical value of post-Archean
terrigenous sediments (Taylor & McLennan, 1985). The variation in the shapes
of REE patterns is partly represented by La/Sm, Yb/Sm and Sm/Nd ratios (Fig.
3.7). REE patterns show that REE from La to Nd are strongly depleted, Sm to
Dy are either depleted of have contents similar to the protolith, and from Ho to Lu
are passively enriched in UHP gneisses relative to medium-LREE metasediments
(Fig. 3.6). UHP gneisses can be divided according to their LREE content into
two groups: ultra-depleted samples with∑
LREE(La − Sm) < 10 ppm and
semi-depleted samples with 10 ppm <∑
LREE(La− Sm) <60 ppm.
Sm/Nd ratios of the non-UHP samples are confined at 0.17–0.24 and close to
Sm/Nd=0.17 of the average continental crust (Taylor & McLennan, 1985) and
Sm/Nd=0.214 in GLOSS (Plank & Langmuir, 1998). In UHP gneisses Sm/Nd
is always higher from 0.23 to 0.86 and there is a weak negative correlation be-
tween LREE and Sm/Nd: ultra-depleted samples and also sample B94-29 have
Sm/Nd > 0.4 and samples with∑
LREE > 20 ppm have Sm/Nd 0.23–0.36.
Th and U concentrations in UHP gneisses are much lower than in non-UHP
samples and Th/U ratios are very variable from 0.14 to 6. Ultra-depleted samples
have very low Th/U ratios 0.14–0.46 while the semi-depleted samples have Th/U
from 2.5 to 6, similar with 4–8 of protolith samples.
Large Ion Lithophile elements: Rb, Cs, Sr, Ba and Pb
Concentrations of Rb, Cs and Sr are very variable but overlapping both in non-
UHP samples and UHP gneisses, with comparable averages between them (Fig.
3.6). The Rb/Sr ratio varies widely, both in UHP gneisses (from 1.4 to 10.8)
and in the potential protolith (Fig. 3.8). In the potential protoliths Rb and
Cs are fairly well correlated (Fig. 3.8). Most of the UHP gneisses occur on the
same Rb–Cs trend as non-UHP samples, but several samples are shifted to lower
Rb/Cs ratios. UHP gneisses show a nice correlation between Rb-K (Fig. 3.6).
Ba concentrations in modern sediments vary widely due to its association
with organic material and hydrothermal precipitates (Plank & Langmuir, 1998).
In the potential protolith metasediments Ba concentrations are within a narrow
range of 240–640 ppm Ba and in the UHP gneisses Ba content is 50–210 ppm and
133
Figure 3.7: Concentrations and ratios of selected trace elements relative to theLREE in the Kokchetav rocks.
9 0.7
8 • 0.6 • • 7 •
0.5 6 • • •
• • 0.4 • . .. 5 • •
:I • "' • • ...J
~4 • • ~ 0.3 • I • 3 • ••• • • 0.2 • 2 ~ • I • UHP gneisses I 0.1 I • UHP gneisses I r • protol ith • • protolith
0 0
0 100 200 300 400 0 100 200 300 400 LREE, ppm LREE, ppm
1200 25 • UHP gneisses
1000 • ~ I • UHP gneisses I • protolith 20 ~ • protolith
800 • • • 15 .. .Q 600 • • z t: i= • ..
400 • ~10 •• ~ rl'• • ._ .. 200 5 •
0 0 0 100 200 300 400 0 LREE, ppm 100 200 300 400
16 LREE m
0.9 I+ I • UHP gneisses I 14 • UHP gneisses
0.8 • protolith • protolith 12 0.7
0.6 10 "0 "' ~0.5 ...J 8
~0.4 ~ :c z 6
0.3 .,.. •• • 4
0.2 ... ... ~. 0.1 2 •
0 0 • •• 0 100 200 300 400 0 100 200 300 400
LREE, ppm LREE, ppm
134
Figure 3.8: Concentrations of LIL elements in UHP gneisses and non-UHPmetasediments.
the average is around 100 ppm. Therefore Ba concentrations in UHP gneisses
are clearly lower relative to protolith (Fig. 3.8). Pb content in UHP gneisses is
0.5-4.7 ppm is generally lower than 1.2-14.5 ppm in non-UHP samples though
there is significant overlap.
High Field Strength Elements: Ti, Nb, Ta, Zr, Hf
Protolith metasediments contain 0.6–0.9 wt.% TiO2 and UHP gneisses tend to
have higher TiO2 (0.8–1.7 wt.%) although there is a significant overlap (Fig.
3.6). Nb and Ta contents in UHP gneisses are generally lower than in non-UHP
samples, but concentrations overlap with the moderate-LREE metapelites (Fig.
3.6). Nb/Ta ratio in non-UHP samples has a narrow range 8.9–11.3 (average 9.9)
and in the UHP gneisses it has a huge variation from 5.7 to 22. In the ultra-
depleted samples the Nb/Ta ratio (19–22) is higher than in the protolith (Fig.
3.7). Sample B94-29 also has a very low LREE content (16 ppm) but oppositely
135
Nb/Ta ratio is as low as 5.7. A chart of Ti/Nb vs LREE shows that with an
increase in LREE the ratio Ti/Nb decreases initially, before showing an increase
at high LREE concentrations (Fig. 3.7).
Most of the UHP gneisses contain 150 to 220 ppm Zr, which is within the
range of Zr content in non-UHP samples (140 to 312 ppm, Fig. 3.6). The only
exception is sample B94-355, which has 77 ppm Zr. The Zr/Hf ratio in the UHP
gneisses is 34.7 ±1 which is indistinguishable from the 35±1 in the non-UHP
samples. Therefore there is no evidence for Zr depletion in the majority of the
UHP gneisses. There is also no correlation of Zr with FeO+MgO and Mn.
Phosphorus
The phosphorus content of the potential protolith metasediments ranges from
230 to 1000 ppm, though the majority of samples fall between 230–400 ppm.
This content is generally lower than the average 1900 ppm P in modern oceanic
sediments (Plank & Langmuir, 1998) and 750 ppm in Proterozoic cratonic shales
(Condie, 1993). In the UHP gneisses the average P content is 400 ppm (190–1080
ppm) thus it is within the range of non-UHP samples. In the UHP gneisses, the
content of P decreases with decrease LREE contents (Fig. 3.9), thus probably P
was depleted from ultra-depleted UHP gneisses with respect to protolith.
Ratios of elements from different groups
The UHP gneisses are strongly depleted in LREE, Th and U, moderately depleted
in P, Ba and Be and the situation is less clear for Ta, Nb and Pb. Therefore
in the UHP gneisses LREE/HREE, LREE/HFSE, LREE/LILE ratios are lower
than in protolith metasediments. However ratios HREE/LILE, HREE/HFSE
and LILE/HFSE are mostly similar to those of the protolith metasediments. For
instance, the important Lu/Hf ratio is 0.07–0.14, with an average 0.10 in the
UHP gneisses, is very close to the Lu/Hf 0.06–0.11, with an average 0.097, in the
non-UHP samples; only sample B94-355, which is depleted in Zr, has a Lu/Hf
0.34.
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Figure 3.9: Concentrations of Be, Pb and P in UHP gneisses and non-UHPsamples relative to other components.
3.6 Discussion: General
3.6.1 Host minerals for REE, Th, U in UHP rocks
One of the key questions for this study is to identify the host for REE, Th
and U in UHP rocks during UHP metamorphism and exhumation. In crustal
granulites, 70–95% of LREE, Th and U are usually concentrated in accessory
phases and the main hosts for LREE and Th are allanite, monazite and other
rare accessory minerals (Bea, 1996; Bingen et al., 1996). K-feldspar can also
host a significant fraction of LREE, although the other major minerals do not
incorporate significant LREE and Th. Zircon is an important host for U. In UHP
eclogites LREE, Th and U are largely concentrated in allanite (Hermann, 2002a).
In the majority of the Kokchetav non-UHP samples, LREE are concentrated
in monazite (section 3.5.2). Chapter 2 presents a detailed description of mon-
137
azite in HP-UHP samples from Barchi Kol’ and demonstrates that monazite was
recrystallized at various metamorphic conditions, and inherited monazite is ab-
sent in these rocks. The UHP gneisses were significantly depleted in LREE,
Th and U relative to protolith (Fig. 3.6) but still minerals of these elements
were identified in several samples. UHP gneisses contain complex associations of
REE, Th and U minerals. As a high pressure LREE phase samples B94-60 and
B118A50 contain allanite, whereas monazite hosts LREE in B94-333. During
retrogression, monazite and allanite decomposed, the Th and LREE were decou-
pled and formed separate minerals (Fig. 3.4). Th became concentrated in Th
orthosilicates (ThSiO4), while LREE were concentrated in secondary monazite.
LREE-bearing minerals were redistributed on the thin section scale and formed
secondary low-Th monazite associated with Ti minerals (monazite near rutile,
titanite in samples B118A50 and B94-60). In sample B118A50 the allanite pseu-
domorph composed of Th-minerals and mica, phengite has high Si content (3.4
Si pfu). This indicates that allanite was replaced at high pressure, because such
phengite is only stable at pressure around 30 kbar (Hermann, 2003).
One can propose two agents that facilitated replacement of REE minerals:
melt, produced by decompression breakdown of phengite, and aqueous fluid.
Monazite and allanite are both stable in granitic melts at crustal and UHP con-
ditions (Hermann & Rubatto, 2009, Chapter 1). Furthermore, melt in equilib-
rium with monazite does not cause large fractionation of LREE relative to Th
(Chapter 1). Therefore the retrogression of HP/UHP monazite/allanite by melts
is unlikely. In sample B94-333 the mineral synchysite (CaCe(CO3)2F ) contains
anions F− and CO2−3 . The presence of these anions in aqueous fluids can sig-
nificantly increase the solubility of REE (Poitrasson et al., 2004; Tropper et al.,
2011). Therefore the more probable scenario is the retrogression of REE miner-
als by fluids. A similar mechanism was proposed for the formation of synchysite
in granites where primary magmatic monazite/allanite is often altered (Forster,
2000).
3.6.2 Geochemical features of the Kokchetav sediments
The protolith for the Kokchetav rocks were shales interbedded with dolomite
limestone sediments and intruded by basic magmas. The sediments were de-
posited at a passive continental margin, while the mafic intrusions occurred dur-
ing continental rifting (Shatsky et al., 1999, 2006a).
Throughout geological history, there have been significant temporal changes
138
in sediment compositions (Taylor & McLennan, 1985). This variation needs to be
taken into account when Kokchetav sediments are compared with modern sedi-
ments. From 1 Ga to 0.5, Ga the development of life on continents caused a grad-
ual change from feldspars-quartz sediments to fine clay-rich sediments (Kennedy
et al., 2006). Changes in mineralogy were accompanied by a decrease in Ca, Fe,
Mn, Mg and an increase in K (Cox et al., 1995). High K2O abundances are a
feature of Proterozoic sediments, differing them from Archean sediments, which
had high Na (Taylor & McLennan, 1985), and modern sediments, which also have
K/Na<1 (Plank & Langmuir, 1998).
Several features of the Kokchetav non-UHP rocks distinguish them from typ-
ical modern sediments and the average continental crust, including their low Sr
and P contents, high K, and high Rb/Cs and K/Na ratios (Table 3.2).
Pelites, greywackes and sands are the main types of clastic sediments. Pelites
and greywackes differ by the grain size and fractionation degree: greywackes
represent the sand fraction, whereas pelites are the finer part of the sediments
(Taylor & McLennan, 1985). Greywackes are formed by physical weathering
and are composed primarily of quartz and feldspars with some fraction of mafic
minerals (pyroxenes, amphiboles), whereas pelites contain high fractions of highly
weathered components — clay minerals. Therefore a high abundance of clay
minerals produces high alumina content and highly peraluminous pelites, whereas
greywackes have compositions close to metaluminous rocks. Quartz sands are
formed by fractionation of clay component from highly weathered rocks and thus
are complimentary to pelites.
Garnet-biotite gneisses are the dominant rock types comprising ≈80 % of the
Kumdy Kol’ and Barchi Kol’ UHP units of the Kokchetav complex (Pechnikov &
Kaminsky, 2008) and often these gneisses contain diamonds. UHP gneisses from
this study are typical garnet-biotite gneisses. Kyanite content is low or negligible
in these rocks.
It is interesting to try to determine the source of the Kokchetav sediments
from geochemical criteria. Plank & Langmuir (1998) proposed discrimination of
sediments sources based on REE patterns. The Yb/Sm ratio of non-UHP samples
is approximately 0.5 and the La/Sm ratio is 5–6. These values are close to the
field of continent derived sediments.
139
3.6.3 Constraining protolith composition
A critical parameter for geochemical modeling of partial melting processes is a
reliable estimate of the protolith composition. It can be obtained from:
a) Average compositions in the literature, such as estimates of the upper con-
tinental crust by Taylor & McLennan (1985); Condie (1993) or GLobal Oceanic
Subducted Sediment (Plank & Langmuir, 1998). The estimate by Condie (1993)
was applied for comparison with the Kokchetav rocks by Shatsky et al. (1999) .
b) Composition of suitable rocks within the same terrain, that have not been
affected by partial melting. Comparison of the Kokchetav non-UHP rocks with
average compositions of the crust shows that many elements have concentrations
typical for crustal sediments but there are a number of significant differences;
in particular the Kokchetav sediments have very high K2O contents and Rb/Cs
ratios (see section 3.6.2). Therefore, the representative average of the Kokchetav
metasediments is more suitable than the average composition of the continental
crust or GLOSS.
When comparing bulk rock compositions of the UHP gneisses and potential
protolith samples, the first observation is the relatively homogeneous composition
of the UHP gneisses and the large variability in composition of the potential
protolith samples. It is thus proposed that the protolith for the UHP gneisses
were a subset of the non-UHP metasediment, rather than a diverse suite.
However, UHP gneisses have higher content of Ca, Na and lower concentration
of Al than non-UHP metasediments of the Kokchetav massif (Fig. 3.5). These
elements should not be much affected by UHP melting because their concentra-
tions in granitic melts are comparable or lower than in Kokchetav UHP gneisses.
Therefore the difference in composition can be attributed only to a difference in
protolith. Non-UHP metasediments from the Kokchetav massif have high alu-
mina contents and correspond to pelitic sediments, however the UHP gneisses
are similar to metagraywakes.
Constraints from melting
Melting of metapelites and metageywackes over a large range of PT conditions
produces melts of similar granitic composition, and “leucocratic components”
SiO2, Al2O3, K2O and Na2O3 constitute >95% of melt composition recalculated
to an anhydrous basis (Montel & Vielzeuf, 1997; Douce & Beard, 1995; Hermann
& Green, 2001). “Mafic components” FeO, MgO, CaO and MnO have low
solubility in granitic melts and behave compatibly, remaining in residue. Melts
140
which coexist with assemblage of quartz/coesite and white mica have peralu-
minous compositions. With increasing pressure, Na2O becomes compatible in
clinopyroxene and K2O contents increase (Hermann & Spandler, 2008).
Fig. 3.5 presents plots of elements constituting granites relative to FeO +
MgO and the trend of bulk rock composition evolution with extraction of granitic
melt in comparison with potential protolith samples. Granitic melts on average
contain 14 wt.% Al2O3 and loss of granitic melt from rocks with lower content
of Al2O3 will cause a decrease in Al content in the restite. This consideration
demonstrates that high-Al protolith samples such as high-Al gneisses from Enbek
Berlyk and metapelite samples B01-3 and CHAG cannot be the protolith of the
UHP gneisses, which mostly contain 10–14 wt.% of Al2O3. Only several samples
of UHP gneisses have sufficiently high Al content to be derived from metapelites.
Orthogneisses are very similar to the granitic melts and have a low content of
mafic components (CaO+FeO+MgO three times lower than in metasediments).
Therefore, extremely high degrees of melt extraction are necessary in order to
produce, from an orthogneiss, a content of mafic components similar to the UHP
gneisses. Such high degrees of melt extraction contradict trace element compo-
sitions and K2O content of UHP gneisses. These considerations are used for the
construction of a self-consistent estimate of Kokchetav Metasediment Composi-
tion (KMC) Table 3.1.
Calculation of protolith composition
The Kokchetav Metasediment Composition (KMC) presented in Table 3.1 was
calculated using the following assumptions:
* For the majority of elements, KMC is based on the averages non-UHP
samples, excluding outliers.
* For SiO2, Al2O3 and K2O concentrations were chosen in order to satisfy
the trends predicted in major element composition (Fig. 3.5).
* For Na2O and CaO the highest concentrations of protolith metasediments
were chosen in order to justify the high abundances of these elements in the UHP
gneisses.
* The protolith rocks form two distinct groups which differ in their LREE
content by a factor of two. Metasediments with high− LREE content of 250–
280 ppm∑
LREE. These samples have high Th (25–30 ppm) and high Th/U
ratios 5–8 (in comparison with Th/U=4 of UCC (Taylor & McLennan, 1985)).
These samples also have an elevated content of Nb (20–25 ppm) and Ta (2.1–2.8
141
ppm). These metasediments have substantially higher trace element concentra-
tions than typical crustal rocks and are unlikely to represent protolith for the
UHP gneisses. Metasediments with moderate− LREE contain 110-140 ppm∑
LREE and 10-17 ppm of Nb and 1.1–1.4 ppm of Ta, which are lower than
in the previous group and close to typical concentrations in sediments (Plank
& Langmuir, 1998). I consider these lower concentrations as representative for
the protolith of the Kokchetav UHP gneisses. An average of 4 metasedimentary
samples with moderate LREE concentrations (B94-256, KY98-3, G03, B95-99)
were taken for REE, Th, U, Ta and Nb estimates for KMC composite.
The resulting average protolith composition is presented in Table 3.1 and
corresponds to high-K metagreywacke with moderate LREE content.
3.7 Discussion: Behaviour of elements during
UHP melting
3.7.1 Major elements
The main feature of major element composition the Kokchetav gneisses is the
negative correlation of K2O and FeO + MgO (Fig. 3.5). This correlation is
due to the variable degree of extraction of a granitic melt from metasediments of
consistent composition. Therefore K2O was incompatible and was depleted, and
FeO, MgO, MnO and CaO were compatible and experienced passive enrichment
in residua. Na2O becomes increasingly compatible with an increase of pressure
of melting (Hermann & Spandler, 2008). However experimental melts buffered
by phengite and omphacite (Hermann & Spandler, 2008) have higher or similar
Na2O content than the Kokchetav gneisses. Thus it is likely that Na2O content
might be reduced or unaffected by melting of UHP gneisses but it cannot increase.
SiO2 and Al2O3 contents in gneisses were not affected by melting. Due to low
content of Fe and Mg in granitic melts melting would not affect Fe/Mg ratio of
restites but increase FeO and MgO contents.
3.7.2 REE, Th and U
A main feature of the UHP gneisses is a strong depletion in LREE, Th and U (Fig.
3.6). Therefore understanding the factors controlling these elements are of crucial
importance for the interpretation of the origin of genesis. LREE, Th and U are
hosted by monazite/allanite at HP-UHP conditions. Monazite and allanite have
142
quite similar solubility and partitioning properties (Klimm et al., 2008, Chapter
1) and for the sake of simplicity in the following discussion monazite will be used.
In the sense of LREE behaviour there are two distinct melting regimes: (1)
with residual monazite present, LREE and Th are controlled by monazite/allanite
solubility, whereas U is controlled by monazite/melt partitioning; (2) complete
dissolution of monazite in melt, then the melt acquires a major fraction of LREE,
Th and U, and other minerals start to control abundances of these elements in
the residue. My experiments (Chapter 1) show that In equilibrium with residual
monazite, U is less compatible than LREE and Th up to at least 50 kbar. In line
with incompatibility of U in monazite during melting a decrease of Th/U ratios
often reported in migmatite granulites (Rudnick et al., 1985; Nozkhin & Turkina,
1993; Schnetger, 1994; Bea & Montero, 1999). Another possible host for LREE,
allanite, also has much lower partitioning coefficients for U than for LREE and Th
(Hermann, 2002a; Klimm et al., 2008). Zircon has strong preference to U relative
to Th (Rubatto & Hermann, 2007a). Therefore Th/U ratios is an important
geochemical indicator for the discrimination of various melting regimes.
There are two distinct groups: ultra-depleted and semi-depleted samples,
which differ substantially in their LREE, Th and U concentrations. In samples
ultra-depleted in LREE, Th/U ratios are much lower than those of the protolith
(Fig. 3.7, 3.10). Ultra-depleted samples can be interpreted as restites after effi-
cient melt extraction when all LREE and Th hosted by monazite were dissolved
in the melt. The control over the Th/U ratio was gained by another phase, most
probably zircon, which was a residual phase and in which U is more compatible
than Th (Rubatto & Hermann, 2007a). Semi-depleted samples preserved a sig-
nificant fraction of their LREE and this can be explained either by the presence
of residual monazite/allanite or the melt extraction from these rocks was less
efficient than in the ultra-depleted samples. Melts buffered by monazite have a
lower Th/U ratio than the bulk rock and extraction of these melts will cause
depletion in U relative to Th and LREE (Chapter 1). However most of the UHP
gneisses have Th/U ratios similar to their protolith, and there are no indications
of depletion of U relative to Th and LREE (Fig. 3.7, 3.10). Therefore it is
unlikely that residual monazite was present during melting and variable LREE
abundances in the UHP gneisses are not related to different degrees of monazite
dissolution. After monazite dissolution in a melt, the melt became the main host
for LREE and the fraction of melt which was not extracted controlled the amount
of REE remaining with the residue and REE patterns or resulting restites.
143
Figure 3.10: LREE, Th and U chondrite normalized patterns of protolith andUHP gneisses separated according their LREE content.
3.7.3 Large Ion Lithophile Elements: Rb, Cs, Sr and Ba
Generally LILE are considered to be incompatible elements during melting pro-
cesses. My data show that only Ba is clearly lower in the UHP gneisses than
in the Kokchetav metasediments. Although Rb, Cs and Sr concentrations vary,
they are broadly comparable with protolith samples and continental crust com-
position (Fig. 3.6). If these elements were depleted then it was to a much lesser
extent than the LREE. Barium is depleted in UHP gneisses but its depletion is
not correlated with LREE content.
The behaviour of the LILE during melting is controlled by feldspars and micas.
In plagioclase, Sr and Ba have high partition coefficients which increase when in
equilibrium with high Al and low Ca melts (Ren et al., 2003) . Alkali feldspar
(oligoclase) has a strong preference for Ba (D≈3–10) and Sr (D≈1–3) relative to
Rb (D≈ 0.3) and Cs (D< 0.1) (White et al., 2003). Sanidine has high partition
coefficients for Ba and Sr, but Rb partitioning is close to unity (Leeman & Phelps,
1981). Phengite in equilibrium with fluid has the highest preference for Rb,
followed by Cs (Melzer & Wunder, 2000), whereas Ba and Sr are incompatible.
In the experiments conducted by (Hermann & Rubatto, 2009), due to the absence
of suitable host phases, Sr behaved as the most incompatible element. However,
in my dataset, Sr does not appear to be depleted in UHP gneisses. This seems
to be in disagreement with experimental studies which observed fractionation of
LILE and strong incompatibility of Sr and Cs (Hermann & Rubatto, 2009). One
can propose several possible explanations for this discrepancy:
144
The characteristic feature of the Kokchetav non-UHP metasediments is the
high abundance of K and Rb, and low content of Sr . Therefore Rb/Sr ratios
of the Kokchetav metasediments are higher than typical crustal values and the
decay of the Rb should have produced highly radiogenic Sr in these rocks.
1) The actual protoliths of the UHP gneisses had a higher Sr content than non-
UHP meta-sedimentary rocks exposed on the surface of the Kokchetav complex.
This is supported by the fact that the Kokchetav metasediments have a much
lower content of Sr (40 ppm) than the average continental crust 350 ppm Sr
(Taylor & McLennan, 1985) or GLOSS 327 ppm (Plank & Langmuir, 1998).
Also, the UHP gneisses have higher Ca content than most potential protolith
metasediments and Sr is the geochemically similar to Ca. However in non-UHP
samples from the Barchi Kol’ unit (B94-333, B95-99) Ca content is comparable
with UHP gneisses but Sr is still very low.
2) A high-Sr component added to the UHP gneisses after melt extraction.
Such a component could be carbonatitic melt, derived from melting the marbles
and carbonate-silicate rocks that are abundant in the Kokchetav complex (Kor-
sakov & Hermann, 2006). Such carbonatitic melts were observed as former melt
inclusions in UHP garnets and pyroxenes and showed a Sr content as high as
1200 ppm (Korsakov & Hermann, 2006). However carbonate-derived Sr is likely
to have a non-radiogenic Sr isotopic composition. On the other hand, weakly
metamorphosed marbles from the Kokchetav complex contain only 35–50 ppm
of Sr (Buslov & Vovna, 2008) and the partition coefficient of Sr between calcite
and carbonatitic melt is close to unity (Ionov & Harmer, 2002).
3) A third possibility might be similar bulk partition coefficients for Rb, Cs,
and Sr during the melting. It seems to be in disagreement with experiments
on melting of metapelites at UHP conditions (Hermann & Rubatto, 2009) which
predicted strong incompatibility of Cs and Sr. However experiments by Hermann
& Rubatto (2009) show a decrease in fractionation of Sr/Rb and Cs/Rb with
temperature and in high T experiments (900oC, 35 kbar and 1000oC, 45 kbar)
ratios for Rb, Cs, Sr are similar to those in starting composition (within a factor
of 2). This can be due to the increase of phengite/melt partitioning coefficients
with temperature.
Shatsky et al. (1999) presented bulk rock compositions of UHP rocks from
the Kumdy Kol’ area of the Kokchetav complex. Shatsky et al. (1999) concluded
that “there is no significant age information in the Rb-Sr data” and that “Sr
isotopes are very radiogenic and are not supported by the present Rb”. A number
of samples in their study have K2O <1.2 % coupled with low Rb, Cs and Sr
145
concentrations. These samples were clearly depleted in K and LILE. It can be
proposed that in the low K2O samples from the study by Shatsky et al. (1999),
effective loss of K2O and Rb was due to complete dissolution of phengite in melt
and efficient melt extraction. Among the rocks presented for this study there
are no samples with compositions analogous to the LILE-depleted samples from
Shatsky et al. (1999).
In summary, it appears that bulk rock compositional data are insufficient to
resolve the discrepancies in LILE content in the Kokchetav UHP gneisses. Some
clues to the fate of the LILE during UHP metamorphism were obtained through
examination of former melt inclusions and mineral compositions (see Chapter 4).
3.7.4 High Field Strength Elements: Zr, Hf Nb and Ta
Zr concentrations are similar in most of the UHP gneisses and in protolith
metasediments (Table 3.1), with the exception of sample B94-355, which has
a particularly low Zr content of 77 ppm. Zr/Hf ratios are also very similar in
UHP gneisses, Kokchetav non-UHP metasediments and in the average continen-
tal crust (Taylor & McLennan, 1985). Therefore Zr and Hf were neither depleted
nor fractionated during UHP melting. Additionally, there is no evidence for the
increase of Zr content in the residue, e.g. there is no correlation between Zr and
Fe+Mg. Therefore, it is likely that the concentration of Zr in the melt and in
the bulk rock were similar. At HP conditions, the main host for Zr and Hf in
metapelites is zircon. The solubility of zircon in a melt show a strong temperature
dependence (Watson & Harrison, 1983) and experiments at 20 kbar by Rubatto
& Hermann (2007a) show decrease of Zr solubility with pressure, but calibra-
tions of Zr solubility to ultrahigh pressure are not yet available. Zr+Hf content3
in the melt buffered by zircon at 1000oC and 45 kbar is 260 ppm (Hermann &
Rubatto, 2009), close to the Zr concentrations in the non-UHP samples and the
UHP gneisses. Extraction of these melts will neither decrease nor increase the
Zr content in the residue but if the temperature was significantly higher than
1000oC, then due to the strong temperature dependence of zircon solubility Zr
would behave as an incompatible element. Therefore Zr behaviour provides ad-
ditional tight constraints that, when melt extraction occurred, temperature did
not exceed 950-1000oC.
Two important geochemical conundrums involve Nb and Ta. The first deals
3Zr and Hf are considered together because experiments by (Hermann & Rubatto, 2009)were doped with Zr and Hf, thus it is likely that zircon in their experiments contained high Hf.
146
with the subchondritic Nb/Ta ratio of the depleted mantle (11–16) and conti-
nental crust (8–14), which is, lower than the Nb/Ta ratio in chondrite meteorites
(18–20) (Rudnick et al., 2000; Jochum et al., 2000; Munker et al., 2003; Nebel
et al., 2010). The “missing Nb” was proposed to be hosted either in the core
(invariably a sink for everything), in eclogites (Rudnick et al., 2000) or in the
Hadean crust subducted to the mantle (Nebel et al., 2010). The second issue re-
gards the pronounced negative Nb and Ta anomaly relative to Th and La present
in the continental crust and in magmas generated in arc settings. This feature is
so characteristic for arc magmas that the Nb–Ta negative anomaly on a spider
diagram is used as a geochemical indicator of discrimination of subduction re-
lated basalts. In experiments on rutile/melt partitioning Drt/meltNb is lower or close
to Drt/meltTa (Schmidt et al., 2004; Xiong et al., 2011; Foley et al., 2000; Hermann
& Rubatto, 2009). Major rock forming minerals such as Pl, Cpx, Opx and Grt
all have very low Dmineral/meltNb and importantly all these minerals as well as rutile
have Dmineral/meltNb /D
mineral/meltTa < 1 (Blundy et al., 1998; Fulmer et al., 2010).
UHP gneisses form two groups according their Nb/Ta ratios. Three ultra-
depleted samples and three semi-depleted samples have Nb/Ta ratios 15–22, sub-
stantially higher than potential protolith samples (Nb/Ta=10) or UCC/GLOSS
(Nb/Ta=11–14). Other samples have Nb/Ta ratios < 10 (lower than the pro-
tolith), and the Nb/Ta is as low as 5.7 in sample B94-29. Whereas low Nb/Ta
ratios can be explained by effect of rutile, rocks with high Nb/Ta ratios experi-
enced fractionation by another mineral.
Tungsten mill potentially can contaminate sample by Ta during sample crush-
ing. However it unlikely to affect measured Nb and Ta concentrations because
samples were crushed in several batches, of which only one batch was crushed in
tungsten carbide mill (which contains trace amount of Ta), the Ta/Nb variation
persist in UHP samples and is absent in non-UHP samples despite the mill used.
Moreover, Ta contamination should cause decrease in the Nb/Ta ratio whereas I
observed a large increase in several samples.
Regardless of the explanation for the observed large fractionation of Nb/Ta,
my data show that restites after melt extraction from metasediments in subduc-
tion zones can be a reservoir for “missing Nb” in the mantle. Also these data
show that melt derived from Kokchetav gneisses should have high LREE and Th
content, but relatively low concentrations of Ta and Nb. Hence, UHP melting
can produce a negative Ta–Nb anomaly in melts derived from metasediments.
147
3.7.5 Phosphorus
Both the non-UHP samples and the UHP gneisses show low and overlapping P
contents. There is, however, a correlation between LREE and P in UHP gneisses
(Fig. 3.9). The ultra-depleted samples have the lowest P contents. Therefore,
P was incompatible during melting and apatite was completely dissolved in the
melt. Indeed, P contents in melts saturated in apatite at 1000oC and 45 kbar,
is 1350 ppm (Hermann & Rubatto, 2009) and in experiments by Green & Adam
(2002) melts buffered by apatite contained as much as ≈1 wt.% of P2O5. There-
fore, even low degree melting would be sufficient to dissolve all apatite in rocks
with 270 ppm P. Ultra-depleted samples contain ≈160 ppm P and this minor
amount of P at UHP conditions is likely hosted by garnet, which can contain
500–2000 ppm of P (Rubatto & Hermann, 2007a).
3.7.6 Sulfur, chalcophile elements
Sulfur content measured by XRF was above detection limit in 4 samples, with
maximum SO3 of 0.9 wt.% in sample B94-57 (Table 3.1). These small number
of samples is insufficient to make a rigorous statistical assessment, but it appears
that there is no correlation of S content with Pb concentration. On the other
hand there are indications of a positive correlation between S and Sb, As and
Zn. The majority of the UHP gneiss rocks have a significantly lower Pb content
(average 2.5 ppm) than the protolith (average 6.6 ppm). Pb is not correlated with
either LREE or with K2O (Fig. 3.9). It is still likely that Pb was incompatible
during UHP melting but factors controlling its behavior are unclear.
Hermann et al. (2006) proposed that sulfide melt immiscibility with silicate
and carbonatitic melt existed at UHP conditions in some of the Kokchetav rocks.
This melt should have a very high density and therefore might sink though the
host rock and scavenge chalcophile elements. Such a process can explain the
positive correlation of S with chalcophile elements. On the other hand high S
contetn can be origianl feature of these rocks.
3.7.7 Be
Average concentrations of Be are similar for both the UHP gneisses (1.76 ppm),
the non-UHP samples (2.4 pmm) and in GLOSS (3 ppm) (Plank & Langmuir,
1998). Be concentration in UHP gneisses shows much larger variations than in the
potential protolith and is not correlated with LREE content. Therefore Be was
148
not controlled by efficiency of melt extraction. Additionally Be in UHP gneisses
is not correlated with K2O, suggesting that phengite and/or melting degree did
not play a role in controlling the Be behaviour. However, samples with low Ca
clearly have reduced Be content (Fig. 3.9) .
Therefore depletion in Be might be unrelated to the melting but rather reflect
fluid loss on the prograde evolution. The amount of Ca controls the presence of
mineral like clinopyroxene, lawsonite, zoisite, amphibole, etc. These minerals
might be the main hosts of Be (Marschall et al., 2007) in metaseiments.
3.8 Discussion: Numerical estimates on melting
of UHP gneisses
3.8.1 Melting model
In this work the terms restite and residue have different meaning, defined as
follows:
• Residue – solid fraction of the rock during partial melting.
• Restite – rock formed after melt extraction and crystallization of residual
melt. Residue consists of resitite and residual melt.
• Residual assemblage – minerals that are present during partial melting.
The general mass balance equations for polyphase system is:
CEltotal =
∑CEl
i ∗ ai (3.1)
1 =∑
ai (3.2)
Where CEltotal is the concentration of element El in the protolith (or bulk
system); CEli concentration in phase i and ai is the fraction of that phase.
Partial melting produces three principal components: solid residue, residual
melt and extracted melt (Fig. 3.11). Residue together with residual melt form
restite after partial melting.
aresidue + aresidual melt + aextracted melt = 1 (3.3)
149
CElprotolith = CEl
residuearesidue + CElmeltaresidual melt + CEl
meltaextracted melt (3.4)
CElrestite = CEl
residuearesidue + CElmeltaresidual melt (3.5)
The solution of this set of equations for quantities of melts and residua re-
quires estimates of the composition of protolith, restite and melt. The protolith
composition was estimated in Section 3.6.3. UHP gneisses represent restites.
Constraints on the composition of the melts and residua for some elements can
be obtained from the experimental data and geochemistry of the UHP gneisses
and they fall in several specific cases:
* An element was highly incompatible and was concentrated in the melt only
(LREE, Th, U).
* An element was compatible and its concentration in the melt can be con-
sidered as negligible (Fe, Mg, Mn, Cr).
* An element was essential component of the melt and its concentration can
be considered as constant (K, Si, Al).
For a very incompatible element, the concentration in residue is negligible.
Therefore, equation 3.4 transforms to:
CElprotolith = CEl
meltaresidual melt + CElmeltaextracted melt (3.6)
On the other hand the fraction of residual melt will control the concentration
of the incompatible element in the restite and its content in the restitic sample
will be:
CElrestite = CEl
meltaresidual melt (3.7)
Combination of the equations 3.6 and 3.7 allows exclusion of melt composition
and simple transformations result in the formula for the calculation of relative
fraction of melt extracted from the rock:
EME =aextracted melt
aresidual melt + aextracted melt
= 1− CElrestite
CElprotolith
(3.8)
This parameter is defined as efficiency of melt extraction (EME) and can vary
from 0 to 1. If EME is low then most of the melt remained with the residue and
if EME is close to 1 then the residue was left almost without a melt after melt
150
Figure 3.11: Schema of system evolution during partial melting. Stage 1: pro-tolith composed of solids. Stage 2: partial melting of rock, forming melt phaseand solid residue. Stage 3: part extraction of the melt, forming extracted meltand the rest of melt forms residual melt. The ratio of extracted melt relative tothe amount of melt is efficiency of melt extraction (EME). Stage 4: crystallizationof residual melt. On cooling the residual melt reacts with residue, crystallizesand forms restite. The ratio of the extracted melt relative to mass of the rock isFraction of Melt Loss (FML). See text for discussion.
extraction.
The fraction of melt loss (FML) can be estimated for elements with known
concentrations in the protolith and melt:
FML = aextracted melt =CEl
protolith − CElresidue
CElmelt − CEl
residue
(3.9)
Parameters EME and FML can be calculated independently for different el-
ements. Values for aresidue, aresidual melt and aextracted melt can be calculated from
3.3.
Modeling of the melting processes is a well developed branch of geochemistry
(Zhou, 2007), though it is generally applied rather to mantle melting than to
crustal rocks. The main types of melting models are:
* Batch melting – melt is generated in one step and then completely separated
from the residue
* Fractional melting – the melt is removed from the source as it is formed
* Dynamic melting – fraction of melt remains in residue, whereas excess of
that fraction is constantly removed
151
* Open system melting – there is addition of external material during melting
The approach presented above is a sort of batch melting model, which takes
into account the variable extent of the melt extraction. The parameters used
above can be directly linked to formulations from other studies. The classic
formula is (Hanson & Langmuir, 1978; Zhou, 2007):
CElprotolith = CEl
extracted melt ∗ F + CElresidue ∗ (1− F ) (3.10)
Where F is melt fraction F = massmelt
massprotolith. In our formalism F = amelt residual+
amelt extracted. Because the residual melt forms various fractions of the restite, the
formula cannot be applied directly to the UHP gneisses. Another interesting pa-
rameter is the fraction of residual melt in the restite FRM = aresidual melt
aresidue+aresidual melt,
which was defined by Zhou (2007) as Ψ– mass porosity of melting residue. Pa-
rameters F, EME, FML and FRM are interconnected as essentially there are
two independent variables. The most sensible representation of these variables
appears as the plot EME vs. F (Fig. 3.14). Below I present calculations of EME
and FML parameters for the Kokchetav gneisses and the behaviour of major and
trace elements is interpreted through these melting parameters.
3.8.2 Estimate of melt loss
The amount of melt extracted for each sample can be calculated by the formula
3.9. If there is a significant difference in element concentration between melt and
residue than this calculation can yield an estimate of fraction of melt loss. SiO2,
Al2O3, Na2O and CaO don’t fit these requirements because their contents in the
granitic melts and in protolith metasediments are similar and because of large
uncertainty in Ca and Na contents in the protolith. The composition of extracted
melt is unknown but experimental studies provide important constraints. Melt-
ing at 1000oC, 45 kbar produces granitic melts with ≈7% K2O and low content
of FeO, MgO, MnO and HREE (Hermann & Spandler, 2008) and extraction
of such melt produces the negative correlation of K2O with compatible elements
in UHP gneisses (Fig. 3.5). From the estimated protolith composition KMC
(Table 3.1, for details of calculation see section 3.8.1) and equation 3.9 fraction
of melt loss can be calculated from concentrations of K2O and FeO + MgO.
These estimates provide similar estimates of FML within 0.1–0.4 and there is
good correlation for estimates obtained from different elements (Fig. 3.12). This
agreement demonstrates that the estimate of the Fraction of Lost Melt is ro-
bust. Averages of FML estimated from K2O and FeO + MgO were used for the
152
Figure 3.12: Fractions of melt loss form residual UHP gneisses estimated from theincrease of FeO+MgO and decrease of K2O show good agreement. Compositionsof bulk rocks are from this study (TS) and from study by Shatsky et al. (1999)(S99).
calculations of FML which are presented in Table 3.3.
3.8.3 Efficiency of melt extraction
The occurrence of ultra-depleted samples and the absence of Th–U fractionation
in semi-depleted samples indicate that, during melting, monazite/allanite were
entirely dissolved in the melt (see section 3.7.2). Therefore LREE, Th and U
behaved as very incompatible elements and during melting were hosted by the
melt only and efficiency of melt extracted (EME) can be calculated by formula
3.8.
The estimate of EME is dependent on the assumption of LREE content in
protolith. The assumed protolith composition provides an estimate of EME from
50 to 98 % for samples from this study (Table 3.3)and those of Shatsky et al.
(1999).
153
3.8.4 Modal abundances of minerals
The current major mineral assemblage in the UHP gneisses is Qtz, Grt, Phe, Bt,
Pl, Kfs and Ky. At UHP conditions, feldspars and biotite were not stable and
over large range of conditions and compositions typical assemblage represented
by Coe, Grt, ±Phe, ±Cpx, ±Ky, ±melt (Hermann & Spandler, 2008).
Phase compositions and reactions for a particular rock composition (pseudo-
section) can be calculated using internally consistent thermodynamic data-bases
(Holland & Powell, 1998) and computational software e.g. THERMOCALC or
PerpleX (Powell et al., 1998; Connolly, 2005). Pseudosection calculations pro-
vide comprehensive information about the mineral assemblage of the rock, but
they are dependent on assumptions regarding unknown parameters like water
content. Also, thermodynamic models for some phases (e.g. melt) are not well
constrained for UHP conditions. In the interest of simplicity, the major mineral
composition of UHP gneisses below are calculated by mass balance of bulk rock
compositions against mineral compositions from experiments (Hermann & Span-
dler, 2008). These calculations are somewhat similar to normative mineralogy
calculation by CIPW algorithm. A first limitation of this approach is that the
melt fraction cannot be calculated together with phengite abundance, because
the melt composition is very similar to Phe + Qtz thus causing internal corre-
lations and inconsistency in the calculation. FeO and MgO have pressure and
temperature dependent substitutions in minerals and for the sake of simplicity
FeO and MgO were replaced by their sum. Results of the calculation are shown
in Fig. 3.13.
3.8.5 Modal abundances of minerals and residual melt
Combined together the EME, FML and mineral abundances provide estimates of
residual melt preserved in residuum. Subtraction of that amount of melt provides
the composition of UHP residue and then the same calculation as above results in
mineral abundances in that residue. The results of this calculation are presented
in Fig. 3.13.
Calculated with these assumptions, coesite mode in residual association is
33–55 %. High abundance of SiO2 ensures that melting was not limited by
available silica. Garnet was an abundant residual phase during UHP melting
(27–63 %). Phengite mode was up to 30 % in samples with low melting degree
to null in intensively melted samples. The modal proportion of omphacite varies
between 0–12 % and depends mostly on the Na content. These calculations reveal
154
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155
Figure 3.14: The effect of melting degree and efficiency of melt extraction onthe rock composition. (a) Principal schema. (b) Plot of EME and F calculatedfor UHP gneisses from this study (diamonds) and from Shatsky et al. (1999)(squares).
that if melt is not accounted, at UHP conditions the Kokchetav UHP gneisses
were composed mostly of garnet and coesite (Grt+Coe 65–98 %), with variable
amounts of phengite and omphacite.
3.8.6 Role of melting parameters on the behaviour of el-
ements
Parameters FML and EME provide complimentary information: FML is an esti-
mate of melt loss relative to residue together with residual melt, while EME de-
scribes the fraction of extracted melt derived from LREE content. UHP gneisses
do not show any correlation between FML and EME (Fig. 3.14 b). This is rea-
sonable: the amount of melt is controlled by parameters like rock composition,
water content, temperature, etc, whereas melt extraction is controlled by the me-
chanical separation of melt. However high F potentially should make extraction
of melt easier.
The samples can be divided into two groups: (1) very depleted samples with
high EME, small FML and phengite present according to the model, and (2)
samples with moderate depletion in LREE, moderate EME, high FML and the
amount of residual melt is sufficient to accommodate all K2O, and thus phengite
was absent. Melt accommodated all of LILE, but because only a fraction of melt
was extracted, there is no strong depletion in LILE in these rocks.
Two samples (256 and 302,5-1) from the study by Shatsky et al. (1999) show
156
both high FML and EME. These samples are characterized by low K2O, Rb, Cs
and also LREE, Th and U contents. I propose that these samples represent a
case when both phengite and monazite were completely dissolved in the melt,
the fraction of the melt was high (in order to accommodate all K2O) and that
melt was efficiently extracted from the residue.
Therefore there are several principal melting regimes:
* low EME: rock does not change composition significantly, no matter how
much melt was present in the system.
* low FML, high EME: phengite is a residual mineral, monazite is dissolved
and efficiently extracted. This is the case for gneisses ultra-depleted in LREE
from this study.
* high FML and EME: all phengite is melted, monazite is dissolved and melt
is efficiently extracted. This is the case for gneisses depleted both in LILE and
in LREE, Th and U from the study by Shatsky et al. (1999).
3.8.7 Melting history of the UHP gneisses
The Kokchetav UHP gneisses might have changed composition during several
principal stages of the rock evolution (Fig. 3.15):
1. at subsolidus conditions, due to loss of aqueous fluid,
2. on the prograde path when the rock crossed the water-present solidus, and
there was loss of low temperature melt,
3. at peak conditions through loss of high temperature melt,
4. during exhumation, by loss of melt produced by phengite breakdown.
These stages might combine in different proportions and the whole picture
appears very complicated. The most extreme scenario is that small melt fractions
were extracted on all the high T part of the PT path. The bulk rock compositions
provide important constraints on the melting history, by limiting the number of
possible scenarios.
High pressure fluids at 600–750oC are very dilute with total solids content
5-10 % (Spandler et al., 2007). The solute in the fluid is mostly composed of
SiO2 with some Al2O3 and Na2O, and is also enriched in LILE relative to REE.
U is also a mobile element in HP fluids and its solubility is increased by oxidising
conditions and the presence of Cl and F in the fluid (Keppler & Wyllie, 1990; Bali
157
Figure 3.15: PT path of UHP rocks of the Kokchetav complex and principalstages when loss of fluid/melt might has occurred. The PT path of UHP rocks isbased on Hermann et al. (2001) with changes according Auzanneau et al. (2006).1 subsolidus loss of fluid, 2 low temperature melting, 3 high temperature, peakmelting, 4 phengite breakdown melting on exhumation. Gray lines show locationof reactions of phengite breakdown, with phengite disappearance right/belowlines: (a) reaction Phe(Ms)+Qtz(Coe)=melt from (Auzanneau et al., 2006), (b)reaction Cpx+Phe+Qtz=Bt+Pl+Grt+melt (Auzanneau et al., 2006), (c) phen-gite upper stability limit (Hermann & Spandler, 2008).
et al., 2011). Also highly acidic/saline fluids can dissolve substantial amounts of
LREE (John et al., 2004; Tropper et al., 2011), but such fluids are unlikely in
the Kokchetav case. The absence of depletion in U, LILE and Na2O in UHP
gneisses makes significant fluid loss (stage 1) seem unlikely.
In the Kokchetav non-UHP metasedimentary rocks monazite or allanite is
the host phase for LREE, Th and U (see section 3.5.2). Monazite and allanite
solubility has a strong temperature dependence and melt formed on the pro-
grade path should be buffered by monazite/allanite (Chapter 1). In equilibrium
with monazite/allanite, U is an incompatible element relative to LREE and Th
(Hermann & Rubatto, 2009, Chapter 1). If melt extraction on the prograde
path was significant then Th/U ratio of the rock should increase. However in
semi-depleted samples, the Th/U ratio is roughly the same as in the protolith.
158
Therefore, significant loss of low temperature melts (stage 2) is unlikely.
Fluid influx can considerably influence the solidus behaviour of the rock. If
a sufficient amount of aqueous fluid is present in the rock then all the phengite
can be dissolved in the melt at relatively low temperatures (∼800oC). However,
the presence of residual phengite in some samples is evidence that there was no
excess water during melting, but rather the amount of water was limited.
The main feature of the UHP gneisses is their depletion in LREE, Th and
U that can occur only if the melt/fluid LREE content was significantly higher
than that of the whole rock. High LREE concentrations estimated in melts which
were extracted from UHP gneisses occur in high temperature melts (Chapter 1).
Therefore the main depletion stage of the UHP gneisses must have happened at
high-temperature conditions (stage 3).
Melt extraction at peak or near peak conditions was incomplete and a sig-
nificant amount of melt remained with residue during exhumation (stage 4).
During exhumation that melt likely reacted with residue minerals and may have
increased in proportion due to decomposition of phengite during decompression.
Subsequent cooling at crustal conditions the melt completely crystallized feldspar,
quartz and new micas.
3.8.8 Nomenclature of the Kokchetav migmatites
UHP gneisses show characteristic depletion in LREE, which can solely be at-
tributed to the extraction of partial melts (see section 3.8.7). Other lithologies
in the Kokchetav complex also show evidence of partial melting, for instance
marbles and garnet-pyroxene rocks contain inclusions of former carbonatitic and
silicate melts (Korsakov & Hermann, 2006). Shatsky et al. (1999) proposed that
trondhjemite veins within eclogites of Kumdy Kol’ UHP terrain were formed by
melting of eclogites. According to the migmatite classification by Sawyer (2008),
the Kokchetav UHP gneisses can by classified as a neosome, because melting
occurred to some extent in most of the lithologies. Similarly, the Kokchetav
UHP gneisses should be called a diatexite, because pre-partial-melting features
are completely erased and replaced by syn-magmatic structures.
Historically in the Kokchetav complex, the term migmatite was applied to the
rocks with well developed melanosomes and leucosomes (Dobretsov & Shatsky,
2004; Ragozin et al., 2009). These rocks are composed of melanosomes rich
in biotite and leucosomes composed of Qtz and Fsp. However, in a modern
terminology, migmatite is defined as any rock which has been partially molten
159
despite its specific textural features (Sawyer, 2008). Therefore all Kokchetav
rocks which show petrographic or geochemical evidence of melting should be
called migmatites.
The study of Shatsky et al. (1999) was based on the 30 samples of various
UHP rocks from Kumdy Kol’ unit of the Kokchetav complex and 6 samples
of non UHP rocks from various parts of the Kokchetav complex. They found
essentially the same characteristics of element depletion as found in this study,
based on the samples from the Barchi Kol’ area. It demonstrates that my samples
are representative and that the loss of LREE, Th, U is not confined to a few
samples of a specific rock type, but the Kokchetav UHP rocks are extensively and
consistently depleted in these elements relative to crustal rocks and Kokchetav
metasediments. Agreement between the two studies indicates that Kumdy Kol’
and Barchi Kol’ have similar melting histories and probably represent parts of a
single UHP block.
The study by Buslov & Vovna (2008) reported compositions of UHP and non-
UHP rocks from the Kokchetav complex and concluded that UHP metamorphism
did not produce any significant effect on their trace element composition, and this
conclusion is in complete disagreement with the present study. This discrepancy
can be explained in several ways: (a) the PT path of the presented samples was
different from the true UHP gneisses, and they either had lower peak conditions or
avoided melting as described in Chapter 2, (b) the selected samples experienced
UHP melting but they had low efficiency of melt extraction, and as only a little
fraction of melt was extracted, anatexis produced insignificant effect on the rocks’
composition (Fig. 3.14). The second explanation seems more plausible and
in any case the discrepancy of the study by Buslov & Vovna (2008) with the
work by Shatsky et al. (1999) and the present study can be explained by biased
sample selection in the study of Buslov & Vovna (2008). It raises interesting
questions about the relative abundance of restitic gneisses versus gneisses with
an unaffected composition. This issue can be resolved by systematic sampling of
gneisses in the Kokchetav complex. Unfortunately only concentrations on major
elements and REE are presented in the study by Buslov & Vovna (2008), thus
their data cannot be included into the dataset for this study.
160
3.8.9 Comparison of Kokchetav UHP gneisses with other
restites
The geochemical and textural features the Kokchetav UHP gneisses are distinct
from crustal migmatites. They are perhaps most easily distinguished by their
geochemistry. In the Kokchetav UHP gneisses, LREE, Th and U are strongly
incompatible elements. The effect of UHP melting on HFSE and on behaviour
of LILE is uncertain. The lower crustal migmatites of the Ivrea-Verbano zone
show significant depletion in Rb, Cs, Bi, Tl, Li and U and moderate depletion
in P, Pb, K and Na (Schnetger, 1994). In other granulite complexes Rb and U
depletion is typicaly reported (Rollinson & Windley, 1980; Nozkhin & Turkina,
1993). However, REE and Th are not usually affected by granulite metamor-
phism (Rollinson & Windley, 1980; Schnetger, 1994). Therefore UHP partial
melting in the Kokchetav had drastically different effects on the trace element
geochemistry of restites when compared to LP melting. Crustal migmatites often
have well defined leucosomes and melanosomes (Sawyer, 2008). In UHP gneisses
such textures are much less apparent.
It has been proposed that specific xenoliths found in kimberlites such as
alkremite (garnet + spinel), corganite (corundum + garnet) and corgaspinite
(corundum + garnet + spinel) were formed after melting and melt extraction
of metasediments at great depth (Mazzone & Haggerty, 1989). These rocks are
extremely enriched in Al2O3, and depleted in SiO2 and alkalis. Ultrapotassic
igneous rocks in the southern Pamir (Tajikistan) contain xenoliths of granulites
and eclogites (Hacker et al., 2005). These rocks originate from the slab subducted
beneath the Asian plate during the collision of India and Asia (Ducea et al., 2003;
Hacker et al., 2005). The metamorphic conditions experienced by these xenoliths
range 875–1100oC and 20–29 kbar (Hacker et al., 2005). The metasedimentary
xenolith with highest peak conditions experienced melting as evident from former
melt inclusions in various minerals (Madyukov et al., 2011). Samples metamor-
phosed at temperature in excess of 900oC show evidence of depletion in La, Th
and Ba (Gordon et al., 2011).
Geochemistry of HP rocks in the Catalina Schist (California, USA) were stud-
ied by Bebout et al. (1999, 1993, 2007). Different units of the Catalina Schist
experienced metamorphism at conditions 275–750oC and 5–12 kbar. Bebout
et al. (1999) studied bulk rock compositions of the Catalina Schist in order to
determine effect of progressive metamorphism on the release of trace elements
and demonstrated that Cs, N, As, Sb and B were extracted from sediments by
161
release of hydrous fluids. Data from this study fail to find a large difference in
the concentration of Cs, Sb and As between non-UHP and UHP rocks. One
explanation can be that all protolith samples in this study are at least medium
grade metamorphic rocks, and thus potentially also lost fluids and thus they are
already depleted in fluid mobile elements. Another explanation can be that PT
path of the Kokchetav UHP rocks was different from that Catalina rocks and
metamorphism occurred at higher pressure. Difference in the behaviour of ele-
ments like As and Sb can be due to reduced conditions in the Kokchetav UHP
rocks (as evident from the common occurrence of graphite and diamond) and
oxidized conditions in the Catalina Schist (though they also contain graphite).
In Chapter 1, on the basis of experimental data, it was concluded that rocks
experiencing melt loss in the presence of residual monazite should acquire de-
pletion in U relative to LREE and Th. Bebout et al. (1999) demonstrated that
Th and U concentrations are similar in rocks of from lawsonite-albite to am-
phibolite facies. Th/U ratios remain constant from lawsonite-albite grade to
epidote-amphibolite, but increase significantly in migmatites of amphibolite fa-
cies. This increase in Th/U can be explained by extraction of U by low degree,
low temperature melts as proposed in Chapter 1.
Rocks of Dabie-Sulu belt in China contain evidence for melting at UHP con-
ditions: there is evidence of addition of granitic melts to eclogites (Zhao et al.,
2007), some eclogites show enrichment/depletion in LREE and LILE (Zhao et al.,
2007) and occasional polyphase inclusions of Kfs+Qtz aggregates are observed
(XiaoYing et al., 2011). Evidence for element mobility by melts are mostly ob-
served in eclogites and not in felsic gneisses. The enrichment of LILE and LREE
in eclogites is, possibly produced by low degree melts of felsic gneisses.
Felsic granulites from the Moldanubian Zone of the Bohemian Massif have
largely undepleted compositions, in particular they have a normal content of
LREE, however concentrations of Th and U are significantly reduced (Janousek
et al., 2004). This geochemical feature was interpreted to be caused by partial
melting of the granulites, and it is considered that redistribution of heat pro-
ducing elements played significant role in exhumation of the HP rocks of the
Bohemian Massif (Lexa et al., 2011).
Castro et al. (2009) proposed that ferrosilicic magmas were formed by near
complete anatexis of greywackes. If this interpretation is correct then ferrosilicic
magmas can by considered as an example of both extremely high melting degree
and efficiency of melt extraction, corresponding to upper right corner of the EME
vs. FML diagram in Fig. 3.14.
162
3.9 Implications
3.9.1 Implications for the exhumation of the Kokchetav
UHP complex
The process of returning deeply subducted rocks to the upper part of the conti-
nental crust is called exhumation, and the details and driving force behind this
process is a long-standing problem in geology. The majority of the theories re-
garding the exhumation of the Kokchetav rocks propose that exhumation is driven
by buoyancy forces, activated by melting of metapelitic gneisses (Hermann et al.,
2001; Dobretsov & Shatsky, 2004). Therefore the density of gneisses at UHP
conditions is a factor of primary importance for models of the exhumation of
Kokchetav UHP rocks.
In the previous part of this chapter I obtained estimates of mineral abun-
dances ±melt (Table 3.3). UHP gneisses were composed of the assemblage Grt,
Coe, Cpx, and Phe (section 3.8.4) and have a density of 3.16–3.3 g/cm3 with
an average of 3.24. Garnet and omphacite are very dense minerals (4 and 3.3
g/cm3 respectively) and they dominate the assemblage. The density of the as-
semblage including melt (assuming a melt density of 2.3 g/cm3 (Connolly, 2005);
this is typical pressure for crustal conditions and is a minimal estimate because
at high pressure density likely higher) is 3.08–3.3 g/cm3 with an average of 3.21
g/cm3. The calculated density of UHP gneisses is thus only marginally lower than
the density of mantle peridotites (3.2–3.3 g/cm3). Therefore, in the scenario of
melt extraction prior to exhumation of UHP restites the buoyancy driven force
seems to be weak. On the other hand, the density of restite minerals together
with residual and extracted melt is 2.8–3.1g/cm3, with an average of 2.97g/cm3,
which is significantly less than peridotite. This indicates an alternative expla-
nation that the exhumation started with melting of gneisses while extraction of
melts occurred only in the later stages of exhumation. These calculation are
in some agreement with calculations by Hacker et al. (2011) that graywackes
metamorphosed at UHP conditions will have relatively low density.
3.9.2 Implications for the trace element signature of sub-
duction zone
The Kokchetav UHP gneisses are apparently the only reported example of restite
whose compositions were significantly changed for many elements by UHP melt-
163
ing during subduction. The main geochemical feature of the UHP gneisses is
their strong depletion in LREE, Th, U and general preservation of HREE, LILE
and HFSE concentrations. Thus melts produced by UHP melting must have
been very rich in LREE, and had strongly fractionated REE patterns. Because
monazite/allanite were completely dissolved in the melts, their Th/U and Th/La
ratios should be similar to those of their protolith (Skora & Blundy, 2010). Melts
produced by UHP melting in the Kokchetav should have moderate enrichment
in LILE and HFSE elements.
The sediment-derived melts formed in subduction zones beneath oceanic arcs
that contribute to the formation of arc magmas are considered to have several
specific features: (1) low Th/U ratios due to the presence of allanite/ monazite
in residue (see Chapter 1), (2) high LILE content with moderate LREE concen-
trations (Hermann & Green, 2001). UHP melts produced from melting of the
Kokchetav gneisses have high Th/U ratios, very high content of LREE and mod-
est content of LILE. Therefore it is unlikely that these melts are direct source of
crustal signature in arc magmas, but rather they may be one of components of the
source region of arc magmas. Nb and Ta are slightly depleted in UHP gneisses.
However Th and U are depleted much more extensively and thus melts extracted
from the Kokchetav gneisses had low Nb/Th and Nb/U ratios, which are typi-
cal for the convergent plate margins (Ryerson & Watson, 1987). Kokchetav UHP
gneisses have Nb/Ta ratios above and below the protolith’s ratio. It demonstrates
that Nb and Ta can be fractionated by UHP melting.
Cosmic radiation in the atmosphere produces 10Be, which has a half life of
1.51 My. Eventually 10Be from the atmosphere resides in oceanic sediments and
is buried during subduction. Fluids and/or melt extract Be from the sediments
(Ryan et al., 1995; Marschall et al., 2007) and the abundances of 10Be in arc
magmas is an important tracer of subduction processes. The presence of 10Be in
arc volcanics was one of the most elegant confirmations of the plate tectonics the-
ory. Better understanding of geochemistry of Be in the rocks from the Kokchetav
complex can provide important independent constraints on the behaviour of Be
in subduction metamorphism. Be is depleted in rocks with low Ca content and
thus probably was controlled by one of the Ca minerals (clinopyroxene, lawsonite
or zoisite). The absence of correlation of Be content with LREE and K2O in the
studied samples might indicate that Be loss did not occur during melting but by
fluid loss during prograde dehydration.
164
3.9.3 Implications for the isotopic signature of enriched
mantle reservoirs
From the point of view of isotopic geochemistry, the Earth is composed of iso-
topic reservoirs. It is widely accepted that the most important isotopic reser-
voirs include primordial and depleted mantle, continental crust, HIMU mantle,
etc. Some basalts sample mantle reservoirs with a signature similar to continen-
tal crust rocks. Usually these reservoirs are linked to subducted crustal mate-
rial, thus it is assumed that crustal material does not significantly change ratios
of principal elements during subduction (Zindler & Hart, 1986; Chauvel et al.,
2008). However Kokchetav UHP gneisses demonstrate that subduction melting
can induce dramatic changes in the composition of crustal rocks. Therefore, one
may speculate on the sort of the isotopic signatures that would develop if the
Kokchetav UHP gneisses would remain in the mantle. This question is particu-
larly interesting because sediments are much more often subducted to the mantle
than they are exhumed back to the crust, which is evident from rare findings of
UHP rocks.
The UHP melting of the Kokchetav metasediments caused a strong decoupling
of LREE from HREE and, in particular, fractionated Sm from Nd (Fig. 3.7).
While Nd partitioned into the melt, Sm remained in garnet, causing > 50 %
increase of Sm/Nd ratio in semi-depleted samples and up to 400 % increase in
ultra depleted samples. Therefore decay of the remaining 147Sm in UHP restites
is expected to produce high 143Nd/144Nd ratios over time.
In contrast to the Sm–Nd system, the Lu–Hf system appears to be relatively
unaffected by UHP melting. The effect of UHP melting on the Rb–Sr system is
somewhat unclear. Shatsky et al. (1999) proposed that many UHP samples are
significantly depleted in Rb relative to Sr, and that the amount of radiogenic Sr
is not supported by the Rb concentrations. In samples from this study, evidence
for systematic Rb–Sr depletion and/or fractionation are absent.
The extraction of Pb from oceanic crust by fluids was proposed by Zindler
& Hart (1986) as an explanation for Pb isotopic signature of the HIMU mantle
isotopic reservoir (mantle with high 238U/204Pb (μ) ratio). In UHP gneisses the
Th, U–Pb system shows complicated relationships. Both Th and U were strongly
incompatible, whereas Pb was a moderately incompatible element. In ultra de-
pleted samples the presence of zircon caused very low Th/U ratios, whereas in
semi-depleted samples Th/U ratios are similar to those of the protolith. Pb de-
pletion is less pronounced than Th and U. Shatsky et al. (1999) concluded that
165
Pb isotope composition in UHP metapelites is unsupported by their current ura-
nium content and provided an age of 1.52 Ga. This is somewhat in agreement
with the incomplete fractionation of Pb from U during UHP melting. Therefore,
Kokchetav type UHP melting can produce a large fractionation of Th from U,
but because Pb is more compatible than Th and U, the ingrowth of radiogenic
Pb will be limited and, even after long residence, the Pb isotopic composition will
probably resemble the common Pb of the continental crust. It is also unlikely
that restites after UHP melting of sediments are responsible for the HIMU man-
tle reservoir. On the other hand some OIB magmas for (instance from Hawaii)
have low Th/U ratios and this feature can be linked to the ultra-depleted restites
observed in this study, where the presence of zircon and complete extraction of
monazite by melt resulted in low Th/U ratios. Hogan & Sinha (1991) demon-
strated that isotopic composition of lead released during crustal melting is con-
trolled by accessory and rock-forming minerals. Low temperature melts might
obtain non-radiogenic lead composition, whereas radiogenic lead will be trapped
in zircon and monazite and will be released only at high temperature. In the
case of Kokchetav melting monazite was dissolved in melt and zircon completely
recrystallized, thus most likely in extracted melts lead had the same isotopic
composition as bulk rocks at that moment of time.
Kokchetav UHP gneisses are still very fertile rocks and have high concentra-
tions of trace elements in comparison with mantle. They can be one of the first
lithologies to melt in the mantle and a small addition of those melts can signifi-
cantly change the isotopic and trace signature of mantle melts. If the Kokchetav
rocks were to be mixed into the mantle they would produce reservoirs with high
incompatible elements content, particularly the LILE and the HFSE. The con-
tent of LREE, Th and U would probably be also be quite high in comparison
with ambient mantle. Such a reservoir would have crust-like Rb–Sr and Lu–
Hf systems but highly radiogenic Nd ratios and changing little with time Pb
isotopic composition. With time, restite after UHP melting would move away
from the main terrestrial evolution line in 143Nd/144Nd vs. 176Hf/177Hf space.
Particularly important is that the Nd isotopic composition could evolve to high
values, which could resemble depleted mantle-like values. Therefore the storage
of Kokchetav-type restites in the mantle will produce a source with quite specific
isotopic characteristics.
166
3.10 Conclusions
The study of bulk rock geochemistry of the Kokchetav gneisses allows us to make
some conclusions about the melting process:
* The significance of the UHP gneisses of the Kokchetav complex is that they
are the only crustal rocks in which partial melting caused large scale depletion
and fractionation of elements.
* The UHP gneisses of the Kokchetav complex were formed from metapelitic
protolith that experienced melting and melt extraction at UHP conditions, which
significantly affected their major and trace element compositions.
* During UHP melting of the Kokchetav rocks LREE, Th and U were strongly
incompatible, whereas LILE, Nb, Ta, Pb and Be were moderately incompatible
elements.
* The depletion of LREE was caused by the loss of granitic melts with high
LREE content. Concentrations of LREE, Th and U in UHP rocks are due to
residual melts in moderately depleted samples, and also due to zircon in ultra-
depleted samples. There was no fractionation of Th from U and LREE from Th
during melt extraction because all monazite/allanite were dissolved in the melt.
* Nb/Ta ratio of the UHP gneisses is significantly fractionated, and whereas in
some samples Nb/Ta ratios are similar to or below protolith ratios, other samples
show significantly increased Nb/Ta ratios.
* The compatible behaviour of other trace elements is due to the low solubility
of the accessory/major minerals which host them.
* For the first time, parameters such as melting degree and efficients of melt
extraction are determined for UHP rocks.
* UHP gneisses have lost large quantities of melt. Calculations predict that
the density of UHP gneisses was similar to mantle rocks or only slightly higher.
Therefore the role of buoyancy in the exhumation of the Kokchetav UHP rocks
is likely to be less significant than previously proposed.
* Melting of metasediments at UHP conditions resulted in large fractionation
of the key parent/daughter trace element ratios. In particular the Sm/Nd ratio
increased and the U/Pb ratio decreased.
167
Table 3.1: Major and trace element compositions of the UHP gneisses from theKokchetav complex. bdl — below detection limit, n.a. — not analysed.
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170
Chapter 4
Polyphase inclusions in an UHP
gneiss
4.1 Introduction
Mineral inclusions in zircon and garnet have proved to be a major sources of
information on the evolution of UHP rocks in the Kokchetav and elsewhere
(Sobolev et al., 1991; Hermann et al., 2001; Katayama & Maruyama, 2009).
Intensive studies have led to the finding of an extensive list of UHP indica-
tors in the Kokchetav: diamonds, high-Si phengite, coesite, high-K pyroxene
and high-Si titanite (Sobolev et al., 1991; Hermann et al., 2001; Katayama &
Maruyama, 2009; Ogasawara et al., 2002). Investigation of inclusions has also
led to the discovery of new minerals kokchetavite (polymorph of KAlSi3O8) and
kumdykolite (polymorph of NaAlSi3O8) (Hwang et al., 2004, 2009) that were
both found as microinclusions in K-rich clinopyroxene. It has also been reported
that Kokchetav rocks contain polyphase inclusions, which are interpreted to rep-
resent crystallised melts. Korsakov & Hermann (2006) found rich associations of
polyphase inclusions in the diamondiferous garnet-diopside rock. The inclusions
were interpreted as carbonatitic melts and two types of silicate melts. Microdia-
monds from Kokchetav gneisses have inclusions of silicate glass with high P and
K contents, which is interpreted to represent former melts (Hwang et al., 2006).
Contrastingly, diamonds from marbles and garnet-pyroxene rocks contain inclu-
sions of hydrous high-K fluids (Hwang et al., 2005). The diamonds also contain
fluid inclusions rich in carbonates and water. Numerous authors conclude that
such fluids and melts played an important role in diamond formation (Hwang
et al., 2005; De Corte et al., 2000; Sitnikova & Shatsky, 2009; Korsakov et al.,
2004).
Polyphase inclusions are also reported in many other UHP complexes. The
Erzgebirge complex in Germany has remarkable rocks with ubiquitous polyphase
171
172
inclusions in garnets, often with diamonds (Stockhert et al., 2001). Metamorphic
diamonds by themselves often contain glassy inclusions with a composition sim-
ilar to granite (Hwang et al., 2006). UHP metapelites from Eastern Greenland
contain polyphase inclusions interpreted as crystallized melts (Lang & Gilotti,
2007). Polyphase inclusions found in garnet and kyanite in rocks from Dora
Maira, Italy, were interpreted as inclusions of supercritical fluids by Ferrando
et al. (2009). In the UHP complexes of Dabie-Shan, China, there are pyroxen-
ites with polyphase inclusions interpreted as fluid precipitates (Malaspina et al.,
2006b).
Chapter 3 presented data on the geochemistry of the Kokchetav UHP gneisses
and demonstrated that UHP metasediments experienced melting and melt ex-
traction, which resulted in strong depletion of UHP gneisses in LREE, Th and
U. The bulk rock compositions provide information on the overall effect of the
UHP melting, but the specific details of this process are unclear. In this chap-
ter, I present the detailed descriptions of the UHP garnet-biotite gneiss from the
Barchi Kol’ that contains abundant polyphase inclusions in garnet. Petrogra-
phy and mineralogy of this sample is investigated and experiments on homoge-
nization of inclusions demonstrate that these inclusions represent former melts.
Compositions of inclusions are used for the reconstruction of the metamorphic
and geochemical evolution of the sample during UHP metamorphism.
4.2 Analytical methods
Phase relations were analysed in polished thin sections via back-scattered electron
(BSE) images on a JEOL 6400 scanning electron microscope (SEM) at the Centre
for Advanced Microscopy, ANU. The phase compositions were determined by
EDS, using an acceleration voltage of 15 kV, a beam current of 1 nA and an
acquisition time of 120 s. Hydrous glass was not stable under the focused electron
beam and a significant decrease in Na2O occurred during acquisition. Analyses
were performed by area scan whenever possible but often, only spot analyses were
possible due to tiny size of inclusions. This problem was overcome by applying a
correction to the measured values. Observations during 2 minutes long acquisition
showed the average loss of Na2O was 20 %±7%. Na concentrations measured in
glass were corrected for this value.
A special sample preparation method was used to ensure the reduction of
contamination and the preservation of garnets, which later were used for the ex-
perimental re-homogenisation of inclusions. The sample was crushed to a grain
173
size of < 3mm with a hydraulic press. Then, a quarter of the sample was pow-
dered in an alumina mill for bulk rock analysis. The rest of the sample was
sieved and garnets 1-0.6 mm in diameter were selected for further experiments.
The fine grained (<0.6 mm) fraction of the sample was used for heavy minerals
separation. The bulk rock composition was measured by the same procedure as
in Chapter 3.
Trace elements in minerals and glasses were analysed using the LA-ICP-MS
facility at RSES, ANU, using a pulsed 193 nm ArF Excimer laser with 100 mJ en-
ergy at a repetition rate of 5 Hz, coupled to an Agilent 7500 quadrupole ICP-MS
(Eggins et al., 1998). Laser sampling of minerals was performed in an He–Ar–H2
atmosphere using a small spot size of 16–37 μm. Data acquisition was performed
by peak hopping in pulse counting mode, acquiring individual intensity data for
each element during each mass spectrometer sweep. Counting was performed for
60 seconds, including a gas background measurement of 20–25 seconds. The LA
data were processed by an Excel spreadsheet created by Charlotte Allen. Mineral
compositions were calculated using NIST-612 (Pearce et al., 1997) as the external
standard and SiO2 content measured by SEM was used as an internal standard
for garnet and tourmaline, Ti for rutile and Ce for monazite. BCR-2G was em-
ployed as secondary standard (Norman et al., 1998). Measurement of K content
was a challenge because K39 has an atomic mass close to that of Ar40, the most
abundant ion in ICP-MS. This results in a K39 high background, far higher than
counts from NIST-612 glass. Hence BCR glass, which has a high content of K2O,
was used as a standard for calculation of K2O content in analysis of inclusions.
4.3 Results
4.3.1 Sample description
Sample B94-26 originates from drill hole 111, in the Barchi Kol’ area, at a depth
of 44.5 m. Major minerals include Grt, Bt, Kfs, Pl and Qtz. The rock is com-
posed bands with different proportions dark (Grt and Bt) and light (Qtz and Fsp)
minerals. The most apparent parts of the sample are melanosome (dominated
by Grt and Bt) and leucosome (mostly composed of Qtz and Fsp, Fig. 4.1 a),
however withing melanosome there are smaller bands with elevated fraction of
Qtz and Fsp and large garnet porphyroblasts. The banding dips at 70o relative
to the vertical direction of the drill core. The melanosome is very rich in garnet
of irregular shape, with bands composed predominantly of garnet with an ag-
174
gregate of biotite and feldspar-quartz between them (Fig. 4.1). Bands enriched
in Qtz-Fsp contain large garnet grains (up to 15 mm, Fig. 4.2) while biotite
rich zones contain smaller grains (<3 mm). The leucosome is dominated by fine-
grained quartz and feldspars and also contains small euhedral grains of garnet.
Bt-Pl symplectite are coarse grained aggregates of biotite flakes and plagiocalse
(Fig. 4.1). Plagioclase in symplectite contains small exolutions of Kfs. Bt-Pl
symplectites are abundant in the garnet-rich bands and are also present around
the large garnet grains (Fig. 4.1 d). Dolomite aggregates were observed in the
matrix of melanosome part of the sample.
Accessory minerals include: Rt, Zrn, Mnz, tourmaline and sulfides (FeS2,
gersdorffite – NiAsS and ZnS). Rutile grains have thin exsolutions of ilmenite.
Small sulfides grains are disseminated though feldspar matrix. Large garnet
grains contain large, rounded mineral inclusions of Ky, Rt, Phe and abundant
polyphase solid inclusions (Fig. 4.1). Kyanite is observed in the sample as
rounded inclusions in garnet or as corroded grains between garnet and Bt-Pl
symplectite (Fig. 4.1 d, 4.2). In melanosome matrix dolomite is present.
Several features differentiate sample B94-26 from many other Grt-Bt gneisses
from Barchi Kol’ unit; the presence of large garnet grains, Bt-Pl symplectites, and
rounded inclusions of rutile and kyanite in garnet. Importantly, garnet contains
abundant polyphase inclusions.
4.3.2 Mineral compositions
Garnet
Garnet composition was studied in thin sections and in garnet separates mounted
in epoxy. Melanosome hosted garnet (Grt-M) have compositions considerably
different from leucosome hosted garnet (Grt-L). The most apparent difference is
higher Ca content in Grt-M (Alm 48–63.5 %, Grs 18–23.5 %, Sps 1–5.7 %, Py
12–26.5 %). Large garnets in melanosome have homogeneous composition and
in some small garnets rims show Mn and Fe increase and Ca and Mg decrease.
Grt-L have lower Ca content (Alm 53–72.2 %, Grs 7–16 %, Sps 0.9–3.5 %, Py
20–28.5 %; Fig. 4.3, Table 4.1). From core to rim, Ca, Mn and Mg increase, and
Fe decreases. The rim compositions in Ggrt-M and Grt-L are similar. Garnet
inclusions in zircon have compositions similar to Grt-M and high Mn content
(Alm 53–57 %, Grs 19 %, Sps 2.4–4 %, Py 21.4–23.4 %).
Na content in garnet is 0.03–0.09 wt.% and is similar in both groups of garnet.
Grt-L have high Y (80–200 ppm), HREE, and P (1000–1500 ppm). Grt-M garnets
175
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Figure 4.1: Sample petrography. a) Scanned images of polished section of sampleB94-26 in directions perpendicular (left) and parallel (right) to the drill core.Dotted line shows the boundary between leucosome (L) and melanosome (M)parts of the sample. b-c) Abundant polyphase inclusions in garnet. d) Kyaniteand phengite mineral inclusions in the large garnet. e) Small garnet grain withcore rich in polyphase inclusions. f) Bt-Pl symplectite consuming kyanite grainadjacent to garnet. g) Tourmaline and rutile crystals in the matrix of the sample.
176
Figure 4.2: BSE image of the large garnet grain. The garnet contains mineralinclusions (Ky, Rt, Phe) and numerous polyphase inclusions. Dashed areas depictparts of the garnet with abundant polyphase inclusions. Bt-Fsp symplectite(Sym) are adjacent to the garnet, which is sitting in a Qtz-Kfsp matrix.
have lower Y (50–100 ppm), HREE, and P (300 or 600 ppm). Zr content in both
Grt-M and Grt-L garnets slightly decreases from 25-30 ppm in cores to 10–15
ppm in rims. In contrast Grt-M garnets have concave REE patterns with higher
LREE content and lower HREE than Grt-L (Fig. 4.3). Grt-L garnets have REE
patterns with flat M-HREE distribution from Sm to Lu and a decreasing trend
for LREE (Fig. 4.3). In cores of large Grt-M garnets, a Eu anomaly is absent but
the majority of both Grt-M and Grt-L garnets have small negative Eu anomalies
(Eu/Eu*=0.5–0.7) in their REE patterns.
Micas and feldspars
Biotite is the only mica present in the matrix of sample B94-26. The biotite
in the matrix contains 2.7–2.8 Si pfu and 3.9–4.5 wt.% of TiO2. Phengite was
exclusively found as inclusions in large garnet porphyroblasts and in zircons.
177
Figure 4.3: Composition of garnet: Grt-M — garnet from melanosome, Grt-L —garnet from leucosome, Zrn inc — inclusions in zircon, Grt-new — garnet crys-tallized in homogenization experiments. a) Triangular Ca-Mg-Fe plot. Arrowsindicate compositional change from core to rim in type-L garnets. b) Diagramof spessartine vs. grossular components in garnet. c) REE patterns of garnetsfrom sample B94-26. d) Compositions of phengite and biotite inclusions in gar-net/zircon and matrix grains.
178
Phengite inclusions in garnet contain 3.1 Si pfu and 1.7–2 wt.% TiO2. Zircons
contain phengite inclusions with 3.20–3.30 Si pfu and 2.66-3.4 wt.% TiO2, and
biotite inclusion with 3.0 Si pfu and 3.9 wt.% TiO2.
In the sample matrix the feldspar minerals are K-feldspar (Ab 6–12%, An
0.5–1%, Or 96-86%) and plagioclase (Ab 78-75%, An 8–22%, Or 2–10%). Kfs
contains exsolutions of plagiocalse and vice versa. Kfs inclusion in zircon has
lower content of anortite component (Ab 9.6%, An 0.3%, Or 90%) than matrix
Kfs.
Rutile
Rutile is present as large (up to 1 mm) matrix grains and inclusions in garnet.
Rutile grains in both textural contexts have ilmenite exolution lamellas. Nb
content in the rutile is 3000-5000 ppm, the Nb/Ta ratio varies from 7 to 22 and
U content is 70-90 ppm (Table 4.1). Zr concentrations in rutile inclusion in garnet
and in matrix grains are similar (790-1090 ppm) and and application of Zr-in-
rutile thermometer (Tomkins et al., 2007) indicates temperatures of 735-770 oC
at a pressure of 10 kbar.
Zircon
The zircons are subdivided into four domains based on their textures and trace
element compositions (Fig. 4.4, 4.5). Domain 1 commonly forms the cores of the
larger grains and has low CL emission with weak zoning (Fig. 4.4). This domain
has low Th/U ratios, high content of HREE and steep REE patterns (Fig. 4.5).
Domain 1 contains 400-500 ppm P, 60-300 ppm of U and 34-72 ppm of Ti. The Ti-
in-Zrn thermometer (Ferry & Watson, 2007) indicates crystallization temperature
of 840-970 oC. Domain 2 forms cores or mantles around domain 1 with high CL
emission and weak concentric zoning. This domain has a similar MREE content
to domain 1 but lower HREE contents. Domain 2 has very high Ti content at
100-300 ppm, for which the Ti-in-Zrn thermometer (Ferry & Watson, 2007) gives
temperature of 1050-1130 oC. Domain 3 forms the mantles and rims on most
grains. It has a CL with high to medium intensity. This domain is characterised
by flat HREE pattern, high but variable Th/U ratios and Ti content of 30-125
ppm (corresponding to a temperature 860-1050 oC). Domain 4 is scarce and
forms rims, it can be recognized by its low CL emission. This domain has REE
patterns similar to domain 3 but has high content of Th and U. Ti content is
as low as ≈10ppm, corresponding to 750 oC. Zircon domains 2 and 3 contain
179
Figure 4.4: CL images of zircons from sample B94-26. Numbers from 1 to 4 markdifferent growth domains of zircons.
inclusions of Phe, Bt, Grt and calcite.
Monazite
Some of the polyphase inclusions contain small (< 3 μm) monazite grains. How-
ever monazite was not found as grains either in the thin sections or within inclu-
sions in garnet. One large grain (> 200μm) of monazite was found in the heavy
minerals concentrate and its composition was determined by LA-ICP-MS. The
monazite grain has high Th (19 wt.%, Th/La 1.2), low HREE content (Yb 5-6
ppm, Y 670-750 ppm) and U (1.4 wt.%, Th/U 14). The concentration of Sr in
the monazite grain is as high as 4300 ppm (Table. 4.1)
Other minerals
Schorl tourmaline forms blue and brown grains with zoned coloration and irreg-
ular shape in the Qtz-Fsp-Bt matrix (Fig. 4.1). Tourmaline has a REE pattern
with enrichment in LREE and positive Eu anomaly (Table. 4.1). Dolomite oc-
curs in the melanosome and has an Fe-rich composition (Ca0.5Mg0.35Fe0.15CO3).
Inclusions of calcite were found in zircon (Ca0.95Mg0.01Fe0.04CO3). A variety of
sulfide minerals were detected (FeS2, gersdorffite - NiAsS and ZnS) as tiny grains
dispersed predominately in the rock-forming feldspars. Also sulfides are present
in the polyphase inclusions.
180
Figure 4.5: Zircon composition. a) REE, Th and U patterns of different do-mains of zircons. b) Histogram of temperatures determined by the Ti-in-zirconthermometer (Ferry & Watson, 2007). c) Th/U vs. temperature.
181
4.3.3 Polyphase inclusions
Polyphase inclusions are abundant in garnet and occasionally found in rutile.
Polyphase inclusions often occupy central parts of small garnet grains (Fig.
4.1e). In large garnet crystals (> 2mm), inclusions concentrate in irregular zones,
spreading from central to outer parts of garnet (Fig. 4.2). Zones rich in polyphase
inclusions also contain rounded inclusions of kyanite and rutile (Fig. 4.1d, 4.2).
Most inclusions have polygonal negative crystal shapes and dimensions varying
from less than 1 μm to 1 mm (Fig. 4.6). Large inclusions have radiating cracks
filled with chlorite (Fig. 4.6 b,d,e). Inclusions smaller than 10 μm rarely have
cracks around them (Fig. 4.6 c). The inclusions contain numerous silicate min-
erals: Qtz, Bt, Phe, Ky, Pl, Kfs, Ba-feldspar (with up to 16.3 wt.% of BaO) and
Chl. Accessory minerals in inclusions are Rt, Ilm, Ap, Zrn, baddeleite, Aln, Mnz,
REE-carbonates and sulfides.
Laser ablation analysis of polyphase inclusions
Measurement of major and trace element compositions of the polyphase inclu-
sions in melanosome garnets was performed by LA-ICP-MS in thick section. The
incomplete ablation is a major problem in reliably estimating of the bulk compo-
sition of the polyphase inclusions. Additionally inclusions exposed on the surface
of thin section are most accessible for analysis but a fraction of these inclusions
are lost during polishing. If a large laser spot size is used and the complete ab-
lation of the inclusions is ensured, then the fraction of the inclusions in the total
mass ablated is small and extrapolation to bulk inclusion composition is unre-
liable. Exposed and subsurface inclusions were ablated together with/without
host garnet. The time resolved ablation patterns showed large compositional
variations during course of the analysis due to presence of minerals with different
compositions. Overall integration of signal averaged out some of these varia-
tions, but compositions of the polyphase inclusion still are highly variable (Fig.
4.7). Variations in concentrations of LREE, Th, U, Zr, Nb and relative abun-
dances are particularly high. Therefore, in order to characterize the composition
of polyphase inclusions and determine whether they can be transformed into a
melt phase, I performed experiments on the homogenization of inclusions at high
P and T.
182
Figure 4.6: Multiphase inclusion in garnet from sample B94-26 (BSE images). a)View of inclusions-rich garnet. Numerous inclusions of various sizes are visible.b) Polyphase inclusion with association Qtz + Rt + Ap + Chl +Mnz. Theinclusion has cracks. c) Two small inclusions without visible cracks. d) Inclusionwith association Qtz + Bt + Pl + baddeleite (ZrO2). This association is unstablebecause baddeleite and quartz should react to form zircon. e) Example of a largeinclusion with numerous cracks filled by chlorite. It contains a number of grainsof Ba-rich feldspar (up to 10 wt.% of BaO). f) Unusual triangular inclusion witha Ba-rich mineral and sulfide grain.
183
Figure 4.7: Primitive Mantle (PM — McDonough & Sun (1995)) normalised traceelement content of polyphase inclusions in comparison with the composition ofthe host garnet (host Grt), which is shown by bold red. Arrows show that contentof an element is below detection limit of LA-ICP-MS.
Experiments on the homogenization of polyphase inclusions
Experiments on the homogenization of fluid and melt inclusions are performed
routinely with volcanic and hydrothermal samples (Roedder, 1984). Recently
homogenization of inclusions also was applied to high grade metamorphic rocks,
but the substantial difference is that, in this case, the homogenization must
be performed at high pressure in order to stabilize minerals hosting inclusions
(Malaspina et al., 2006b; Bartoli et al., 2011). Works on fluid and melt inclusions
often aim at estimating conditions at which inclusions formed. In particular the
temperature of inclusions homogenisation, with relevant corrections, can be a
reliable estimate of the temperature of their formation. Instead, in this work,
the main purpose of experiments was the transformation of heterogeneous asso-
ciations into the homogeneous glass suited for analysis by micro beam methods.
Another aim is to demonstrate that inclusions indeed represent trapped melt.
Such experiments should be performed at conditions where garnet is stable, the
hydrous granitic melt is above the solidus and temperature is high enough for
complete dissolution of secondary accessory minerals. These requirements dic-
tate the choice of experimental setup. The experiments were performed at high
184
temperature (900-1000 oC) in order to homogenize inclusions and to dissolve com-
pletely the accessory minerals. High pressure is necessary for the stabilization of
garnet. The experiments were performed with in a powdered matrix medium in
order to cushion garnet grains from the strain occurring in the press assembly. In
order to prevent leakage of the fluid from the inclusions most of the experiments
were performed with an addition of Al(OH)3, which releases water on heating.
High water activity at the conditions used in the experiment presents a problem
because it facilitates the formation of matrix melt and destabilizes garnet. As an
attempt to cope with this effect some of experiments were performed with oxalic
acid (H2C2O4), which produces H2O and CO2 upon ignition and results in low
activity of H2O in fluid (Table 4.2).
Experiments on the homogenisation of polyphase inclusions were performed
on a piston-cylinder press at RSES, ANU. The experiments were run in � 3.5
mm gold capsules at temperatures 900 and 1000 oC and pressure of 10-20 kbar.
Several experiments were performed with garnets separated from the crushed
sample (Table 4.2). The capsule was filled with layers of garnet grains separated
by layers of matrix powder (Fig. 4.8). In the experiment C3388 garnet grains
were set in a mix of 90% SiO2 and 10% Al(OH)3. A one hour long run at 1000 oC
and 20 kbar produced melt, Qtz, Ky and two types of garnet. Approximately
10 % of garnet after the experiment was represented by homogeneous grains
with small glass inclusions and had composition similar to type-L garnet. In an
optical microscope, inclusions appear transparent and isotropic without bubbles
or crystalline phases (Fig. 4.8). An exception was the presence of a tiny grain
of gersdorffite (NiAsS) in one glassy inclusion (Fig. 4.8). The other ≈90 % of
garnet recrystallized and obtained spongy textures after the experiment. The
spongy garnet had lower Ca content and higher Mg# than the original garnet
(Fig. 4.3). The matrix was composed of quartz and melt with variable major
composition similar to andesite: SiO2=60%, Al2O3=13% (Table 4.3). Kyanite
was present in melt near garnets.
Experiment D1348 was performed with a large garnet grain (5× 2.4× 2 mm)
cut from the sample melanosome (Fig. 4.1 a) with a diamond saw. The garnet
grain was placed into the capsule and immersed into the matrix of SiO2+Al(OH)3.
In this run large fraction the garnet survived the experiment and glassy inclusions
were present in garnet, though most of the inclusions have thin cracks. In this
experiment garnet had type-M composition both before and after experiment.
Additional experiments were performed at various conditions (Table 4.2), but
were unsuccessful, generally due to decomposition of garnet. In order to reduce
185
Grt new
Grt FeNiAsS
inclusions
SiO2melt
inclusionGrt
a
Ky
Grt new
100 μm
d e
100 μm
b c
100 μm 100 μm
glass
Figure 4.8: Experiment for the homogenisation of polyphase inclusions. a)Scheme of experimental charge. Garnet layers in the capsule were interlayered bya mix of SiO2 −Al(OH)3. b) Optical image of polyphase solid inclusions beforeexperiment. c) Homogeneous inclusions after an experiment at 1000 oC, in thesystem Grt+SiO2+Al(OH)3. (d-e) Experimental products contain two types ofgarnet: garnet (Grt) with homogenized inclusions and spongy garnet (Grt new)recrystallized during experiment.
melting during experiment, I also performed experiments at a lower temperature
(900 oC) and with reduced water vapor activity/pressure. This was achieved by
running experiment D1300 with mix of SiO2 with oxalic acid (H2C2O4). Inclu-
sions in these experiments homogenized partially, with some silicate crystals (Bt,
Pl, etc.) still present.
Experiments with different set up produced inclusions in garnets with differ-
ent compositions. Homogenized inclusions in garnet from melanosome (Grt-M)
produced in experiment D1348 will be denoted as type-M inclusions and inclu-
sions in Grt-L from experiments C3388 and D1300 will be marked as type-L
inclusions. In two experiments with homogeneous glass composition (C3388 and
D1348) major element composition of glasses was determined by EDS and trace
186
element composition by LA-ICP-MS.
Major element composition of the homogenized inclusions
In general, homogenized inclusions are felsic melts with oxide totals ≈90%. These
melt totals indicate high water content in melts. Melts have 58-68 wt.% SiO2
and 6-10 wt.% of FeO (Table 4.3). Their compositions on a TAS diagram are
close to syenite/andesite. Experimental melts coexisting with garnet and mica at
30-45 kbar conditions have MgO and FeO contents < 1.3 wt.% at temperatures
≤ 1000oC (Hermann & Spandler, 2008). The high content of FeO and MgO in
homogenized inclusions is interpreted to result from interaction of melt with the
host garnet and for small inclusions host garnet also might make contribution.
Fortunately both host garnet and garnet dissolved in inclusion have the same
composition and thus can be treated as one component. A value of 0.8 wt.
% FeO was chosen as original iron content in the melt (Hermann & Spandler,
2008). Then fraction of garnet component in the inclusions was calculated from
equation 4.1 and garnet component subtracted from analyses. Such calculations
show that typical inclusions contain 10-25 % of garnet component. Uncorrected
and corrected compositions of inclusions normalized to anhydrous 100 % are
presented on Table 4.3 and Fig. 4.9.
After subtraction of the garnet component, inclusions have broadly granitic
composition with 2-4 wt.% of Na2O and 2-5 wt.% of K2O (Fig. 4.9). Average
SiO2 content is 73.4 wt.%4 in type-L and 71 wt.% in type-M inclusions, which is
in excellent agreement with experimental melts which contain 68.7-74.9 wt.% of
SiO2 (Hermann & Spandler, 2008). There are some differences in major element
compositions of type-M and -L inclusions. Type-L inclusions are peraluminous
granitic melts and have ASI from 0.8 to 1.7 and average 1.2. Type-M inclusions
have lower content of Na + K and higher Ca and Al (ASI is also close to 1.2)
than experimental melts. These particularities can be due to the crystallization
of garnet of different composition on inclusion walls and thus lead to an incorrect
subtraction of a “garnet component”. Another possible explanation is partial loss
of alkalies through fractures which are visible in garnets before and after exper-
iments. Type-L inclusions have a significant Cl content of 0.4-1 wt.%, whereas
in type-M inclusions Cl content is close to limit of detection of EDS microprobe
(<0.3 wt.%) (Fig. 4.9).
4Here compositions are recalculated to anhydrous basis
187
Figure 4.9: Major element compositions of glassy inclusions obtained after ho-mogenization experiments. a) Na2O vs. K2O. b) SiO2 vs. Na2O + K2O. c)Na2O + K2O vs Cl.
Trace element composition of homogenized inclusions
Homogenized polyphase inclusions often are very small (3-10 μm, rare ≈15μm)
and their LA-ICP-MS analysis represent mixes of garnet with glass thus a calcula-
tion of an absolute concentration is necessary. The compositions of the inclusions
can be calculated if there is an element with an independently constrained con-
centration in the inclusions — an internal standard — CElstdinc (Halter et al., 2002).
The mass fraction of an inclusion in the ablation analysis – X can be calculated
from the concentration of the internal standard element by the formula:
X =CElstd
mix − CElstdhost
CC
Elstdinc −Elstd
host
(4.1)
Where CElhost is the concentration of element ‘El‘ in the host mineral, CEl
mix is
the measured concentration and CElinc is the inferred concentration in the inclusion.
Then, for all other elements, the concentration in the inclusion can be calculated
using the formula:
CElinc = CEl
host −CEl
host − CElmix
X(4.2)
According to this equation uncertainty in internal standard will propagate
188
to uncertainty in the concentration of all elements and thus the choice of the
internal standard needs careful consideration. Elements compatible in garnet in
garnet cannot be used as internal standards because their decrease in garnet-
inclusion mixes is covered by noise of signal from ablation with small spot size.
Na, K and Rb are the elements which are most suitable as internal standard
because they have a low content in garnet and their concentrations in melts can
be estimated from SEM analyses and/or experimental data. Concentrations of
Na and K measured in inclusions vary significantly and the primary reason for
this is the difficulty of microprobe measurement of volatile elements in micron
sized inclusions of different shape and size. Though corrections have been ap-
plied to the Na content in the inclusions it is possible that analytical uncertainty
was not accounted for completely. On the other hand experimental studies are
quite consistent regarding granitic composition of HP-UHP melts derived from
metapelites, both Na2O and K2O vary significantly depending on melting con-
ditions and bulk system composition (Hermann & Spandler, 2008), but Na2O
+ K2O is close to 8 wt.%. This value is very close to Na2O + K2O of type-L
inclusions. In type-M inclusions Na2O + K2O is systematically lower (average
5.6 wt. %) than in experiments. According to experimental data by (Hermann &
Rubatto, 2009) K/Rb ratio in melts buffered by phengite is 10-30% higher than
K/Rb ratio in the starting composition. Therefore Rb content in the inclusions
can be estimated by applying a coefficient to K content in the inclusions, but Rb
has the advantage of being measured by LA-ICP-MS more reliably than K. The
Rb content in the inclusions was estimated at 250 ppm from the K/Rb ratio of
bulk rocks and the K content of the inclusions.
Therefore Na, K and Rb may be used as an internal standard either of Na, K
and Rb, either through the measured concentrations in the inclusions, or using
the values predicted by experiments. Here the compositions of the inclusions were
recalculated to the average concentrations of Na2O and K2O in type-M and -L
inclusions measured by SEM and was extrapolated to Na2O + K2O content of
8 wt.%, which is an average of the compositions of the melts from the study by
Hermann & Spandler (2008). Then X was calculated as average of estimates
from Na, K and Rb. The fractions of inclusions in the LA-ICP-MS analyses
(X) estimated using these constraints are in good agreement between each other
(Fig. 4.10). This is considered as the most robust and consistent approach and
inclusions composition calculated this way is reported in Fig. 4.12 and in Table
4.4. Another way is to calculate for Na2O and K2O content measured by SEM,
which predicts a higher fraction of inclusion in the analyses and thus lower trace
189
Figure 4.10: Estimates of fractions of homogenized inclusions in ablated materialfrom concentrations of Rb and K, in comparison with estimates obtained fromNa concentrations. Estimates from different elements agree between each otherclose to 1:1, thus demonstrating internal consistency. See the text for details.
Figure 4.11: Trace element composition of homogenized inclusions. Chondrite(McDonough & Sun, 1995) normalised patterns of measured garnet-inclusionmixes and composition of the host garnet (host Grt). Arrows show elementswhich concentrations in garnet are below detection limit.
elements concentrations.
Elements compatible in garnets (in particular HREE) cannot be reliably esti-
mated in inclusions because two reasons: firstly uncertainty on garnet component
is secondly, even if concentrations of these elements would be estimated they still
will be completely useless because high garnet capacity for these elements will
result in easy and rapid exchange with host garnet either during exhumation or
190
Figure 4.12: Average compositions of two types of homogenized inclusions andcomposition of protolith and sample B94-26.
during experiments. Therefore only elements highly and moderately incompat-
ible in garnet will be considered in further discussion. LA-ICP-MS analysis of
the homogenized inclusions revealed increased concentrations relative to the host
garnet for many elements: Na, K, As, Rb, Sr, Zr, Nb, Cs, Ba, La, Ce, Pr, Nd,
Th and U (Fig. 4.11). Compositions of the homogenized inclusions are much
more consistent than those of polyphase inclusions, but there remains a large
variability which can be explained only by variable inclusion compositions.
In both type-M and type-L inclusions have similar concentrations of Na, K,
Rb and Cs. Ba content in type-M inclusions is around 1000 ppm and in type-L
inclusions it is highly variable (200-4000 ppm). Sr content is higher in type-M
(500-700 ppm) than in type-L inclusions (30-130 ppm) . There are significant
differences in concentrations of other elements, in particular LREE, Th, U, Zr,
Nb, As. Type-M inclusions have high LREE content (>100 ppm La) whereas
type-L inclusions have a lower LREE content (< 50 ppm La). Also type-M
inclusions have high concentration of Th and U and Th/U ratio of ≈6, and
type-L inclusions have high U content but low Th, and thus low Th/U ratios
of 0.2-0.8. Type-M inclusions have much lower Zr content (30-100 ppm) than
type-L inclusions (230-1000 ppm). Type-L inclusions have high As content (100-
1000 ppm in type-L compared to 30-130 in type-M), which is also confirmed by
observation of gersdorffite - NiAsS in those inclusions (Fig. 4.8).
191
4.4 Discussion
4.4.1 Origin of polyphase inclusions
There are several mechanisms for the formation of polyphase inclusions in min-
erals: 1) capture of melt and its later crystallization; 2) capture of mineral ag-
gregates during growth; 3) capture of melt together with crystal(s); 4) capture
of mineral inclusion and its later decomposition with formation of melt (Perchuk
et al., 2005) or polyphase aggregate.
Several features indicate that the polyphase inclusions represent trapped
melts, rather than mineral aggregates captured during garnet growth. (1) Most
inclusions can be homogenized to a single melt phase. (2) After the subtraction
of dissolved garnet, the inclusions have the composition of peraluminous granitic
melt, similar to high pressure melts produced in experiments (Hermann & Span-
dler, 2008). (3) Inclusions have significant concentrations of incompatible trace
elements (Zr, LREE, Th, U), which are mostly insoluble in major minerals, but
can have high concentration in the melt. (4) Inclusions formed by the capture
of mineral aggregates or melt together with crystal(s) should have wide varia-
tions in composition corresponding to the abundances of phases in the original
inclusion. However, inclusions have a reproducible composition (Fig. 4.9, 4.11),
which cannot be assigned to any simple mineral mixture.
In crustal rocks, crystallization of melt inclusions sometimes produces sim-
ple mineralogy and textures similar to that of granites - nanogranites (Cesare
et al., 2009). Instead, polyphase inclusions in sample B94-26 have a rich as-
sociation of minerals (over 15 phases ) and some components are present in
different minerals: Zr is found in zircon and baddeleite; LREE are present in
monazite/allanite/LREE-carbonates, Ti is found in rutile or ilmenite. Some of
the associations present in the inclusions cannot be in equilibrium, for instance
quartz associated with baddeleite in some inclusions (Fig. 4.6) is particularly
good evidence for absence of equilibrium, because these minerals should react to
form zircon. Thus, variable mineral associations in inclusions are evidence for
the complex crystallization history of polyphase inclusions in UHP gneiss. One
reason for such variability can be decrepitation and loss of fluid phase at various
PT conditions during exhumation.
Composition of inclusions can be modified by several processes, in particular
by partial loss of melt/fluid through cracks, and loss or intake to/from host min-
eral by diffusion. In sample B94-26, rows of very small (less then 1 μm) inclusions
192
of chlorite are commonly present near the large polyphase inclusions (Fig. 4.6,
4.8). Some LA-ICP-MS trace element analyses of garnets are enriched in Rb, Ba
and K. This could be explained by the contamination of the garnet analyses by
chlorite inclusions. At the same time, garnet with chlorite inclusions have the
same content of LREE and Th as non-contaminated garnet. LREE and Th can be
transported rather via melts, whereas LILE can also be liberated by aqueous flu-
ids (Spandler et al., 2007). Therefore the low LREE content in chlorite inclusions
indicates that they were formed from aqueous fluids originated by decrepitation
of inclusions, from fluid exsolution during crystallization of melt. However, small
inclusions usually do not have cracks and only small inclusions survive homoge-
nization experiments. Therefore I assume that homogenized polyphase inclusions
are likely to preserve their original composition. Diffusion can significantly affect
composition of polyphase inclusions, especially in the case of inclusions hosted
by garnet, which has high content for several trace elements. However, content of
trace elements like LILE, LREE and some of HFSE in garnet is very low. Even
if diffusion of these elements was fast enough, but low capacity of garnet would
prevent from efficient transport from inclusions to outer space. On the other
hand change of PT conditions will result decreased solubility of some elements in
garnet and their diffusion out of grain or exsolution to lamellas and inclusions.
Melt inclusions can become sink for the elements which solubility in garnet de-
creased. Hence elements with high concentrations in garnet can be easily affected
by diffusion and thus should considered with caution.
Stoeckhert et al. (2009) proposed that polyphase inclusions with diamonds
in rocks from Erzgebirge, Germany, experienced loss of fluid during exhumation.
Stoeckhert et al. (2009) also concluded that decrepitation of inclusions is con-
trolled by the rate of decompression and that in order to achieve brittle failure of
the host garnet, the decompression had to be extremely fast. If decompression is
not fast enough then inclusions should expand by plastic deformation of garnet.
On this basis Stoeckhert et al. (2009) concluded that exhumation was very fast.
Chlorite filled fractures around some of polyphase inclusions in the Kokchetav
gneiss B94-26 are similar with those of the Erzgebirge inclusions. Peak PT pa-
rameters and exhumation path were similar in The Kokchetav and in Erzgebirge.
Hence, the conclusions of Stoeckhert et al. (2009) that the exhumation was very
fast are likely applicable to the Kokchetav gneiss as well. This conclusion is in
agreement with short upper limit on exhumation time estimated by U-Pb dating
in previous works (Hermann et al., 2001) and in this study in Chapter 2.
193
4.4.2 Interpretation of homogenization experiments
In experiment C3388 homogenized inclusions composed of glass with high con-
tent of K2O, Na2O and Cl (Table 4.3, Fig. 4.9). Matrix melt and secondary
inclusions in the spongy garnet have low quantity of these components . This
difference in the composition of inclusions and matrix melt demonstrates the iso-
lation of inclusions from matrix melt because Na and K are able to diffuse rapidly
in melt (Acosta-Vigil et al., 2006). In experiment D1348 some fraction of bi-
otite/feldspars were intergrown with large garnet grain used as starting material.
In experiment D1348 matrix melt obtained elevated Na and K concentrations
from biotite and feldspars which were together with garnet porphyroblast and
the difference with inclusions and matrix melt is less apparent.
Experiments C3388 and D1300, performed with garnet separates, produced
homogeneous and/or partly homogeneous inclusions in garnets. The garnet
grains with inclusions had compositions (low content (2-5 %) of CaO), which are
equivalent to leucosome garnet cores (Grt-L) and are significantly different from
the garnets occurring in melanosome (Grt-M with 7-8 % CaO). The inclusions
encountered in these experiments correspond to type-L, with low LREE content.
Only in experiment D1348 (with a large garnet grain cut from the sample) has
high-Ca (Grt-M) garnet survived the experiment and this experiment produced
type-M inclusions with high LREE concentrations. Therefore my homogeniza-
tion experiments C3388 and D1300 resulted in efficient selection of garnets and
inclusions: Grt-M garnets with high CaO content were not stable during the ex-
periment and recrystallized completely; only Grt-L garnet with low CaO content
survived. Experiments C3388 and D1300 give a biased result, indicating that
only low LREE inclusions are present in the sample. This study of the UHP
gneiss sample demonstrates that performing only one type of experiment with
either separated grains or with single large garnet grain can give biased result
and careful, extensive petrographic study of the sample is absolutely necessary
for correct interpretation of homogenization experiments.
4.4.3 Estimation of temperatures of inclusions formation
LREE content of melts buffered by monazite and or allanite has strong temper-
ature dependence and can be used as geothermometer (Montel (1993), see also
Chapter 1). The application of the monazite solubility thermometer from Chap-
ter 1 requires estimates of several parameters: LREE and H2O content in the
melt, pressure and α - activity of LREE in monazite. For the protolith of the
194
Kokchetav UHP gneiss, α is ≈0.94 (Chapter 3 and Table 3.1). Type-M inclu-
sions likely formed at peak conditions and thus pressure can be estimated as 50
kbar. The content of LREE is 780-1160 ppm and the calculated temperature is
980-1020 oC. This is a minimum temperature estimate because it is possible that
LREE concentration in the melt was not buffered by monazite. Type-L inclusions
occur in retrogressed leucosome and pressure of their formation can be estimated
at 20 kbar. LREE content in these inclusions is 40-220 ppm corresponding to low
temperatures of 600-750 oC. These temperatures are also minimum estimate be-
cause monazite may have not been saturated during formation of these inclusions.
The uncertainty in the calculation of the absolute composition of the inclusions
from the mixes data propagate an uncertainty in calculated temperatures. If the
composition of the type-M inclusions is estimated from measured concentration
of Na2O and K2O then temperatures are only slightly lower at 930-980 oC.
4.4.4 LREE, Th and U evolution during melting
Sample B94-26 is severely (80%) depleted in LREE, Th and U relative to sed-
imentary protolith (see Chapter 3, section 3.8.1, Tables 3.1, 3.2 and 3.3). The
Th/U ratio of the bulk rock B94-26 is 3, slightly lower than Th/U ratios 4-8
observed in the protolith (Table 3.2). This weak enrichment of U relative Th can
be due to the effect of residual zircon. Type-M inclusions have Th/U ratios of
7±3, which are within range of Th/U ratios of the protolith. These high Th/U
ratios of type-M inclusions indicates their formation at a time when monazite
was completely dissolved in the melt because in the presence of residual mon-
azite/allanite melts acquire low Th/U ratios (Chapter 1). Type-M inclusions
have very high LREE content ≈8 times higher than estimated for the protolith
(KMC, Table 3.1). Therefore, the extraction of type-M melt is likely responsible
for the depletion of the gneiss in LREE, Th and U (Fig. 4.13).
Type-L inclusions have low LREE contents and low Th/U ratios between 0.3-
1, with an average of 0.5. The occurrence of these inclusions in leucosome garnet
indicates their formation during the exhumation and cooling stage. A possible
mechanism can be the formation of additional melt by decomposition of phengite
on decompression, which diluted the residual high-LREE melt (recorded in type-
M inclusions) that remained in the rock after peak melting and melt extraction.
Then type-L inclusions were formed when monazite saturation was reached again
and monazite consumed Th and caused the enrichment of the melt in U. The low
Th/U ratios of these inclusions can be linked with domain 4 of zircons which also
195
Figure 4.13: PT diagram for sample for sample B94-26 with PT path based onHermann et al. (2001) with changes according Auzanneau et al. (2006). Ellipseswith numbers from 1 to 4 denote stages of zircon growth (zircon domains 1–4) andred polygon is for type M inclusions and blue polygon for type L inclusions. Graylines show location of reactions of phengite breakdown, with phengite disappear-ance right/below lines: (a) reaction Phe(Ms)+Qtz(Coe)=melt from (Auzanneauet al., 2006), (b) reaction Cpx+Phe+Qtz=Bt+Pl+Grt+melt (Auzanneau et al.,2006), (c) phengite upper stability limit (Hermann & Spandler, 2008).
have low Th/U ratios and were formed at low temperatures. Another possibility
is that, together with melt inclusions, garnet captured small zircon grains with
high U content, causing high Zr contents in type-L inclusions.
Type-M inclusions have Th/La ratios of 0.8. Type-L inclusions tend to have
high Th/La ratios with an average of 2. Thus the monazite grain from the heavy
mineral concentrate which has very high Th content and Th/La ratio of 1.2, can
be interpreted as having formed by crystallization of type-L melt in leucosome.
The origin of these local variations is unclear, however, they seem to play a minor
role in bulk Th/La fractionation because the Th/La ratio of sample B94-26 is 0.3,
within the range of 0.3-0.6 in protolith. Therefore, the bulk rock composition is
in agreement with insignificant fractionation of Th from La during UHP melting
as proposed in Chapter 1.
196
4.4.5 Hosts for LILE and their behaviour during melting
Section 3.7.3 demonstrated from bulk geochemistry data that the behaviour of
LILE was controlled by the degree of melting and melt extraction, and overall
LILE are weakly depleted in the majority of UHP gneisses. In sample B94-
26, LILE are now hosted in micas and feldspars. Phengite is absent as a matrix
mineral in sample B94-26 and found only as inclusions in garnet and zircon. LILE
concentrations were determined by LA-ICP-MS in matrix plagoclase, K-feldspar
and biotite and also in phengite and biotite inclusions in zircon and in phengite
inclusions in the large garnet porphyroblast (Table 4.1, Fig. 4.14). In the latter
case, inclusions were ablated together with the host zircon. The trace element
composition of micas was reliably estimated by using Al as internal standard,
because zircon has negligible LILE contents, Al and K. The results are presented
on the Fig. 4.14. There is a remarkable similarity between LILE contents between
sample B94-26 and in the protolith for the Kokchetav UHP gneisses (Fig. 4.14
a; see Chapter 3 for protolith estimate). K-feldspar, plagioclase and biotite have
different concentrations of LILE: Pl hosts Sr, Kfs concentrates Rb, Sr and Ba,
Bt has high content of Rb, Cs and Ba (Fig. 4.14 b). Altogether, these minerals
are capable of hosting all LILE present in the rock.
High pressure mica inclusions in garnet and zircon have a higher content of Sr
and Ba than matrix biotite and 3-5 times higher concentrations of Rb, Cs and Ba
than the bulk rock composition (Fig. 4.14 c). Thus, the presence of 20-30 % mica
of this composition would easily accommodate all the Rb, Cs and Ba present in
the rock. The concentration of Sr in mica inclusions in zircon is lower than in the
bulk rock and mica inclusions in garnet have Sr comparable to the bulk rock. In
inclusions of type-M and -L, concentrations of LILE are quite similar (Fig. 4.14
d) with the exception of Sr: in type-L inclusions Sr is lower than the bulk rock
content, whereas in type-M inclusions Sr is 2-3 times higher.
Phengite and biotite inclusions were observed in domain 3 of zircons, and
close to zircon rims. Thus these micas were formed during decompression, but
a high content of Si and Ti indicates their HP-HT origin. Compositions of mi-
cas and feldspars show that in the current assemblage, LILE are fractionated
between Pl, Kfs and Bt, but at HP conditions they were mostly concentrated in
phengite. Type-M inclusions formed at peak conditions have only slightly higher
concentration of Rb, Cs and Ba than the bulk rock and protolith, thus extraction
of this melt would not produce a large depletion in these elements. Sr content
in inclusions is higher than in the rock and protolith, thus it is possible that
197
Figure 4.14: LILE concentrations in the bulk rock and in protolith (a) and com-parison with: b) LILE concentrations in rock forming minerals, c) LILE con-centration in the high pressure mica inclusions in garnet and zircon, d) LILEconcentration in homogenized inclusions.
real protolith had high Sr content. Sr was incompatible and possibly similar Sr
concentration in restite and protolith occur only by coincidence.
4.4.6 Hosts for HFSE and their behaviour during melting
In sample B94-26 concentration of HFS elements (Ti, Nb, Ta, Zr and Hf) are
similar to protolith, suggesting moderately incompatible behaviour during UHP
melting (section 3.7.4).
Sample B94-26 contains 6500 ppm Ti. Ti content in type-L inclusions is 1000-
3000 ppm and 3000-6000 ppm in type-M inclusions. Concentrations in type-M
inclusions are in agreement with Ti solubility of 2800-5500 in 1050-900 oC UHP
melts buffered by rutile (Hermann & Rubatto, 2009). However, these concen-
trations should be considered with caution because garnet can have significant
Ti content and will modify Ti content in inclusions. Nb content is 26-40 ppm
in type-M inclusions and it is < 10 ppm (averaging 3 ppm) in type-L inclusions
198
whereas in bulk sample B94-26, Nb content is 15 ppm. It is expected that type-
M inclusions have Nb/Ta ratio <10, and the extraction of that melt caused an
increase in the Nb/Ta ratio from ≈10 in protolith to 13 in restitic sample B94-26.
However Ta was not measured in the inclusions due to expected low concentra-
tions. Overall, the extraction of type-M melt had very little effect on Ti content,
but it reduced Nb content in the rock and presumably depleted Ta more strongly
than Nb.
Sample B94-26 has elevated Nb/Ta ratio of 13, relative to Nb/Ta ratio of
10 in the protolith. Rutile contains 2800-5200 ppm Nb and 140-600 ppm of Ta.
Nb/Ta ratios in rutile are highly variable between 7-22. In garnet Nb and Ta
concentrations are very low (<0.1 ppm). Phengite and biotite have high Nb
content (10 ppm Nb, Table 4.1) and Nb/Ta ratio in both phengite and biotite
is around 30, more than two times higher than bulk rock. The fractionation of
Nb/Ta ratio by phengite is huge, in particular when compared with rutile, which
has Drt/meltNb /D
rt/meltTa 0.4-1 and is close to unity in low-T granitic melts (Schmidt
et al., 2004; Xiong et al., 2011). Therefore, the increased Nb/Ta ratio in sample
B94-26 can be explained only by the Nb/Ta fractionation in the residual phengite
(see also discussion in section 3.7.4). Though rutile was present in sample B94-26
it is clear that rutile was not the only major host for Nb and Ta in sample B94-26
and it is phengite which has made large positive effect on Nb/Ta ratio. Large
variation in Nb content and Nb/Ta ratios in rutile in sample B94-26 is probably
related to Ti exsolution from HP garnet and reactions of rutile with mica.
Zr and Hf bulk rock concentrations in Kokchetav gneisses are mostly unaf-
fected by UHP melting (section 3.7.4). Type-M inclusions have low Zr content
(30-80 ppm) and type-L inclusions have high Zr content (500-700 ppm). These
concentrations are in complete disagreement with inferred high temperatures for
type-M inclusions and low temperature for type-L inclusions as zircon saturated
melt at 1000 oC and 45 kbar contains ≈260 ppm Zr (Hermann & Rubatto, 2009).
Also low Zr content is in disagreement with the common observation of Zr bear-
ing minerals in polyphase inclusions (zircon and baddeleite). Zr content increases
in garnet with increasing temperature and pressure, and experiments by Rubatto
& Hermann (2007a) estimated very high solubility (200-800 ppm Zr in garnet),
though these data need to be confirmed. In the garnets from sample B94-26 Zr
content is only 10-20 ppm. Therefore one explanation for the high/low Zr con-
centrations in inclusions can be diffusion of Zr between garnet and its inclusions.
Another explanation can be the heterogeneous capture of small zircon grains
together with melt. The difficulties in interpretation of Zr concentration in inclu-
199
sions in comparison with LREE are due to high Zr versus low LREE content in
garnet. It demonstrates that elements incompatible in the host preserved much
better in inclusions. The interpretation of LREE, Th and U concentrations in
inclusions was made possible from the experiments presented in Chapter 1.
4.4.7 Hosts for Be and its behaviour during melting
Sample B94-26 has low Be content - 0.9 ppm, almost 3 times lower than 2.4 ppm
in the protolith. It indicates depletion of Be in sample B94-26 and is consistent
with the high Be content of 10 ppm measured in type-M inclusions. Be content
in HP phengite inclusion in garnet is 2.2 ppm (similar to bulk rock content in
protolith), suggesting that while phengite is a significant host for Be but still
insufficient to impose a compatible behaviour on Be in the bulk rock.
4.4.8 Numerical estimates of melt loss
The bulk rock composition of sample B94-26 is typical for the Kokchetav UHP
gneisses. Relative to the protolith, sample B94-26 has a decreased content of K2O
and increase in MgO and FeO. Using the same calculations for the estimate of
melt loss from UHP sample as in Chapter 3 and taking KMC composition as
protolith, one can calculate that sample B94-26 had a melting degree of 40 %,
and had lost approximately 33% of its original mass as granitic melt. With
efficiensy of melt extraction at 0.8 only 20 % of melt remained in the rock and
at the peak conditions the residue was composed of 10 %Phe, 40 % Coe and
roughly 50 % Grt. According to the classification from Chapter 3, sample B94-
26 is semi-depleted restite, because its LREE content is above 10 ppm, the Th/U
ratio is close to the ratio of non-restitic metasediments and Th/U ratio of the
rock was controlled by residual melt rather than by residual zircon (Table 3.1,
section 3.6.3, 4.15).
On the other hand the degree of melting can be estimated from the com-
position of the inclusions. Type-M inclusions represent high temperature melt
which formed when all of the monazite/allanite was dissolved in the melt. As
I have estimates of LREE 120 ppm in protolith and inclusions provide ≈1000
ppm LREE in the peak melt, the melt fraction can be calculated at ≈10 %
which is significantly lower than the 33 % melt loss estimated from bulk rock
data. An additional complication is that sample B94-26 experienced the loss of
least two different melts and it contains some fraction of leucosome. Two stages
200
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Figure 4.15: Mineral sequence diagram for B94-26. 1, 2, 3 and 4 are differentzircon domains. o – stage when mineral was homogenized by diffusion. ? — itis uncertain whether melt was present in the rock on the prograde path. M —inclusions in melanosome garnet, L — inclusions in leucosome garnet.
of melt extraction and fraction of leucosome in the sample can explain the dif-
ference between melt abundance estimated from inclusions and from bulk rock
composition.
Constraints on behaviour of LREE, Th and U from trace element com-
position of zircon
Zircon is a key accessory mineral in these crustal rocks as it is stable during the
entire metamorphic evolution and is a common restitic phase after melt extrac-
tion. Zircon contains significant amounts of trace elements such as REE, Th
and U that are strongly affected by melting at UHP conditions. Additionally,
zircon growth zones can be precisely dated and the Ti-in-zircon thermometer can
provide temperature estimates.
Zircons from sample B94-26 have four growth zones with distinct texture and
trace element compositions (Fig. 4.5). (1) Zircon cores are dark in cathodo-
luminescence (CL), have low Th/U ratios, high HREE contents and steep REE
patterns. The low Th/U ratio is attributed to formation in equilibrium with mon-
201
azite, which hosted a large fraction of the bulk Th. The steep, HREE-enriched
patterns indicate zircon equilibrium with high-HREE garnet, which formed dur-
ing prograde metamorphism when the garnet fraction was low, and garnet had
high-HREE content (see Chapter 2). Ti-in-zircon thermometry indicates crystal-
lization temperature of 875-970oC (Fig. 4.13).
(2) Zircon mantles show weak concentric zoning and are bright in CL. These
zones have a similar content of MREE to the cores but lower HREE contents,
indicating that they formed when more garnet was present. The very high Ti
content of 100-300 ppm indicates a temperature of formation of 1020-1170oC.
The Th/U ratios are not fractionated with respect to the bulk rock, indicating
that monazite was not present during exhumation. This can be explained by
the complete dissolution of monazite in a melt formed at peak metamorphic
conditions.
(3) Zircon mantles and rims have high to medium CL emission and are char-
acterized by low-flat HREE patterns, and high but variable Th/U ratios. These
domains formed during slight cooling as they have similar Ti-in-zircon temper-
atures as the cores (850-1025oC). The low U content and Sm/Nd ratio of this
zircon domain reflects formation after the melt extraction that produced LREE,
Th and U depletion.
(4) Zircon rims with very weak CL emission have REE patterns similar to
domain 3 but higher Th and U contents. The Ti content is as low as 10 ppm,
corresponding to 750oC, indicating that these zircons formed during retrograde
melt crystallization, when monazite saturation was achieved again.
In zircons from sample B94-26 there is a decrease in HREE content from do-
main 1 to domain 3 and domains 3 and 4 have similar HREE content. These
variations in HREE content in zircon are inevitably connected with garnet, be-
cause garnet is the major host for HREE. The garnet fraction typically increases
with increase in pressure and temperature in HP metasediments (Hermann &
Green, 2001; Auzanneau et al., 2006). HREE is almost entirely hosted by garnet
and increase of garnet fraction will result in decrease of HREE content in gar-
net and consequently in zircon. However in zircon the decrease in HREE from
domain 1 to domains 3-4 is is about 10 fold and it is hard to account for such
a decrease by a change in garnet fraction along. One explanation can be that
changes in zircon/garnet partitioning also played a significant role in changes in
zircon composition. Temperature increase causes a decrease of DZrn/GrtHREE (Ru-
batto & Hermann, 2007a), and it may be that pressure also makes effect on
zircon/garnet partitioning. Another explanation can be that zircon composition
202
reflects equilibrium partitioning with the outer rims of garnet, which can differ
significantly from the compositions of other garnet domains (samples B01-3 and
B94-333 in Chapter 2).
Eu anomaly in minerals (Grt, Zrn and Mnz, etc) is often applied as an impor-
tant tool for the identification of the presence of feldspar in an assemblage and
thus for a discrimination of HP minerals from low pressure associations (Rubatto
& Hermann, 2003, 2007b). The application of this approach to the Kokchetav
rocks from this study is difficult because in general there are no large variations
in the Eu anomaly in minerals. In particular different zones of garnets and zircon
from sample B94-26 have very modest variations in Eu*. Grt, Mnz and Zrn have
Eu* around 0.6, similar to the bulk rock, and Pl, Kfs and Tur have a significant
positive Eu anomaly (Table 4.1). This is enigmatic because textural evidence
(Fig. 4.1, 4.2) suggests that, Tur, Mnz, and external zones of Grt and Zrn very
likely have formed at low pressure, in equilibrium with feldspars and thus should
have a stronger negative Eu anomaly than the UHP minerals. For instance Grt-
L occurs in feldspar rich leucosome and contains inclusions formed at ≈700oC,
however Eu anomaly in Grt-L is essentially indistinguishable from melanosome
Grt-M. A potential explanation can be that Eu was in Eu3+ state during meta-
morphism. The presence of positive Eu anomaly in plagioclase precludes this
interpretation. Another explanation is that in the case of gneiss B94-26 the
main host for Eu was garnet, and thus plagioclase did not produce significant Eu
anomaly in coexisting minerals. It also may be possible that Grt-L have grown
before feldspar in the same leucosome.
4.4.9 Comparison with other data on fluid/melt inclu-
sions from other UHP rocks
Glass and polyphase inclusions have been were observed in diamonds from UHP
rocks in several works (Hwang et al., 2006). The occurrence of these inclusions in
diamonds provides direct evidence of their formation at UHP conditions. How-
ever the tiny size of inclusions in diamonds and the applied method (Transmitted
Electron Microscopy) only allowed semi-quantitative estimates on the major el-
ement composition and no trace element data.
Hwang et al. (2006) reported observations of polyphase, glass and fluid inclu-
sions in diamonds from Kokchetav gneisses and presented semi-quantitative data
about their compositions. The major components of the inclusions are Si, Al and
K, with low content of Ca, Fe and Mg. Therefore the major element composition
203
of the inclusions are comparable directly to melts produced in this study. In di-
amonds from Kokchetav marbles and garnet-pyroxene rocks Hwang et al. (2005)
described inclusions of COH-potassic fluids with up to 40 wt.% K2O. These rocks
are possibly formed by the reaction of granitic melts derived from metasediments
with metacabonates and the difference in bulk system composition may explain
the presence of different type of fluid than those observed in this study.
Melt/fluid inclusions were reported in Kokchetav rocks in several previous
studies (Hwang et al., 2005; Korsakov & Hermann, 2006; Hwang et al., 2006).
Hwang et al. (2006) found silicate glass inclusions in diamonds from UHP Grt-
Bt gneisses. The reported semi-quantitative compositions of these inclusions are
characterized by high contents of K, P, and Cl and by low contents of Al, Na, Fe,
Mg and Ca. The compositions are quite similar to the granites, but they differ
from the inclusions reported in this study by the very high contents of P and Cl,
and low Al. Korsakov & Hermann (2006) described polyphase inclusions in UHP
minerals of the diamond rich calc-silicate rocks from the Kumdy Kol’ area of the
Kokchetav complex and determined their major and trace element compositions.
They observed three types of inclusions: 1) polyphase carbonate inclusions; 2) sil-
icate Ttn-Ep-Bt trace element-rich inclusions (Ce 40 ppm,∑
LREE ≈90 ppm);
3) silicate Kfs-Bt inclusions with low content of REE (∑
LREE ≈0.5 ppm).
The first and second types of inclusions were explained as two immiscible melts
coexisting at UHP conditions. The silicate polyphase inclusions with low LREE
content requires an additional explanation. These inclusions have such a low
content of LREE that they were formed either at very low temperature (while at
low temperature fluid but not melt is present) or they were never in equilibrium
with LREE bearing phases. One can propose that the low-LREE inclusions from
the study by Korsakov & Hermann (2006) can form by decompression melting of
association Phe+Qtz. In this scenario, the association Phe+Qtz appears above
the melting curve and forms melt, whereas phengite only inclusions have a higher
melting temperature and survive. Melting of inclusions bearing association Phe-
Qtz will produce melt without contact with accessory phases and thus it will
have low trace element contents.
In the study by Korsakov & Hermann (2006), Ttn-Epi-Bt silicate inclusions
have Th/U ratios of 0.4. This is comparable with Th/U of 0.3-0.7 in type-L in-
clusions from sample B94-26. These Th/U ratios are different to the typical ratio
of 3-6 observed in the majority of crustal rocks, in the semi-depleted samples
of Kokchetav UHP gneisses and in the high-LREE type-M inclusions from this
study. Ttn-Ep-Bt inclusions described as having “high trace element” actually
204
have a moderate content of LREE (Ce 40 ppm,∑
LREE is 90 ppm). These
values are very similar to the LREE content in type-L inclusions estimated in
this study, thus it is proposed that the diamodiferous marbles were formed by
the metasomatism of carbonates by a melt similar to type-L inclusions. Type-L
inclusions were formed during the exhumation of the UHP gneiss, but it is likely
that prograde, low-T melts should have similar characteristics: moderate content
of LREE and low Th/U ratios. One possible explanation is that Ttn-Epi-Bt sil-
icate inclusions actually represent prograde melts. However the inclusions were
observed in Cpx formed at peak UHP conditions with lamella of K-minerals.
Another explanation can be that high-LREE melts derived from metasediments
experienced fractionation of a phase with high LREE and Th contents, for in-
stance titanite (Prowatke & Klemme, 2005).
Sulfides and arsenides were observed in the rock matrix, polyphase inclusions
and homogenization experiment products (Fig. 4.8). This is particularly inter-
esting because Ni and As content in Kokchetav rocks is very low (114 and 5 ppm
respectively in sample B94-26). Similarly NiAs was reported in inclusions in dia-
monds from Erzgebirge (Hwang et al., 2003) and the mineral Ni, (Fe, Ni, Co)9S8
occurs in diamonds from Kokchetav (Hwang et al., 2001). The association of this
relatively rare mineral with UHP diamonds was interpreted by Hwang et al.
(2003) as dissolution of sulfides in melts and their catalytic effect on diamond
nucleation. There is no depletion of Kokchetav gneisses in Ni and As (Chapter
2), therefore their incompatible behaviour is unlikely. Instead it is possible that
sulfides formed immiscible droplets in silicate melt.
Metaperidotites metasomatised by crust derived fluids are found in China and
sometimes the garnet in them contains hydrous polyphase inclusions. Malaspina
et al. (2006b,a, 2009) performed experimental homogenization of inclusions and
determined their trace element composition. Though absolute concentration in
inclusions have not been determined (due to absence of suitable internal standard)
the trace element signature of inclusions was obtained. The main characteristics
of inclusions are high concentration of LILE, low Th/U ratios and relatively low
LREE concentrations. The characteristics of fluids from the study by Malaspina
et al. (2006b) can be explained by its formation at the presence of monazite
which is able to fractionate Th from U (Chapter 1). The low LREE content
also indicates that the complete dissolution of monazite was not achieved in the
source of the fluid which metasomatized peridotites. According to this inter-
pretation, data presented by Malaspina et al. (2006b) demonstrated that laser
ablation of polyphase and re-homogenized fluid inclusions can provide very sim-
205
ilar results. This is in marked contrast with the enormous scatter of analyses of
polyphase inclusions in this study (Fig. 4.7) and the consistent composition of
re-homogenized inclusions. This difference is because in this study I encountered
high temperature crustal melts with extreme concentrations of trace elements.
Therefore re-homogenization of polyphase inclusions is absolutely necessary for
such high grade rocks, whereas for lower temperature fluids laser ablation of
non-homogenized inclusions probably can provide satisfactory result.
Perchuk et al. (2008) performed experiments where garnets from middle tem-
perature eclogites (15-32 kbar, 600-680 oC) were subjected to high pressure high
temperature conditions of 40 kbar and 800-1100 oC. Before experiments, the gar-
net contained mineral inclusions of clinozoisite, quartz and rutile. After high-T
high -P treatment clinozoisite inclusions decomposed and formed melt patches
which reacted with host garnet and changed its composition. These experiments
demonstrated that melt inclusions can form, not only from melt itself, but also
by decomposition of previous mineral inclusions. Indeed such mechanism is po-
tentially possible, but granitic composition and high trace element abundances
of inclusions, described in this study, preclude such interpretation.
Madyukov et al. (2011) presented study of former melt inclusions in granulite
xenoliths from Pamir, which originate from the same suite of deep crustal xenolith
derived from recently subducted slab (Ducea et al., 2003; Hacker et al., 2005;
Gordon et al., 2011). Former fluid and melt inclusions were observed in Qtz,
Cpx, Grt, Pl, Ap, Tnt and scapolite — all minerals composing xenolith. Former
melt inclusions are composed of a fluid bubble of dense CO2 fluid and a glass with
low water content (1-4 wt. %) and significant content of CO2 and Cl. Inclusions
have highly fractionated REE patterns and have HREE depletion relative to
bulk rock (HREE can be measured in this case because Qtz and scapolite have
low content of HREE). LREE content in inclusions in quartz and scapolite is
300-500 ppm for which monazite solubility thermometer from Chapter 1 gives
temperatures of 830-870 oC. These temperatures are somewhat lower than the
1000 oC estimated by Madyukov et al. (2011), possibly due to undersaturation of
melt in monazite. The Th/La ratios of the inclusions are 0.3, which is a typical
value for metasediments, and thus indicates the absence of any significant Th/La
fractionation in agreement with conclusions of Chapter 1.
In this study high Th/U ratios of high-T inclusions were used as an evidence
that they formed when monazite was completely dissolved in the melt. Melt in-
clusions in enclaves of crustal metapelites from El Hoyazo, Spain have low Th/U
ratios and contain monazite as residual mineral (Acosta-Vigil et al., 2010), which
206
agrees with my interpretation. Interestingly matrix glass in enclaves and partic-
ularly in dacites have much lower Th/U ratios, possibly because monazite was
completely dissolved in the high-T melts. For an inclusion in scapolite Madyukov
et al. (2011) have reported Th and U concentrations and high Th/U ratio of 7.8
(Madyukov et al., 2011). This high Th/U ratio is close to that observed in this
study and it can be interpreted as evidence for complete dissolution of monazite
during melting. Therefore my new data and interpretation of previous works
demonstrate that Th/U ratios of granitic melts are a sensitive indicator of the
presence of monazite/allanite in equilibrium with the melt.
The melting of the Kokchetav gneisses produced granitic melts with high
concentrations of LREE, Th and U, which are very different from typical crustal
granites which contain in the order 100 ppm of LREE. Obviously it raises a ques-
tion about the occurrence of such unusual melts in nature. In the Bohemian
massif there has been reported the finding of polyphase inclusions in spinel peri-
dotites composed of Phl + calcite (Cc) + apatite (Ap) + graphite (Grp) + rutile
(Rt) + monazite (Mnz) + thorianite and some other unidentified phases (Nae-
mura et al., 2008). Dating of thorianite from these inclusions gave an age of 334
±4 Ma, close to the age of the HP metamorphism in the adjacent gneisses. These
polyphase inclusions can be considered as analogous to the inclusions described
in this chapter and they can be interpreted as the result of interaction with the
mantle and the crystallization of high-LREE melts produced by Kokchetav type
melting of metasediments. This study together with the findings of Naemura
et al. (2008) demonstrate that high-LREE melts play significant role in element
transport from subducted slab to the mantle.
4.5 Conclusions
This work presents one of the first studies of polyphase inclusions and their ex-
perimental re-homogenization. The conducted experiments and analytical work
allow to conclude the following :
# High temperature melt inclusions in crustal rocks most likely have high
concentrations of trace elements and thus contain accessory minerals. This makes
such inclusions unsuitable for bulk ablation and homogenization by high pressure
experiments must be performed before chemical analyses.
# Rock can contain several generations of mineral with inclusions of differ-
ent compositions. Different mineral compositions can behave in a very different
ways even during short re-homogenization experiments and these experiments
207
can induce a bias in the selection of inclusions. A thorough examination of
pre-experimental and post-experimental minerals is necessary and detailed pet-
rographic description and mapping are absolutely essential.
# Elements with high concentrations in host mineral (Zr, Ti, HREE in garnet)
can easily change their concentrations in inclusions and thus cannot be used as
reliable petrological indicators. Only elements with low concentrations in the
host mineral such as LREE, Th, U, LILE (K, Rb, Cs, Ba, Sr), Nb, Ta and Be in
the case of garnet can be preserved in high T inclusions.
For the studied UHP gneiss the high pressure homogenization experiments
proved unequivocally, that polyphase inclusions represent former hydrous granitic
melts. During the exhumation occurred partial decrepitation of inclusions which
resulted in variable and non-equilibrium mineral associations in them. Inclusions
record two types of melt, each formed at different stages of rock’s evolution:
* Inclusions with high LREE content (type-M) have high Th/U ratios. These
inclusions represent melt formed at peak conditions when monazite was com-
pletely dissolved in the melt.
* Inclusions with low LREE content (type-L) were formed during the retro-
gression in the leucosomes from residual low-T melt.
Melting occurred at the presence of residual association of Grt, Coe, Phe,
Zrn and Rt. LREE, Th and U were very incompatible. K, Rb and Cs resided
in phengite, and were only slightly incompatible during melting. Sr and Ba were
concentrated by mica in lesser extent and possibly they were more incompatible
than Rb and Cs. Nb was incompatible during melting. The anatexis of metased-
iments produced strongly fractionated granitic melts with high concentrations of
LREE, which were extracted and became effective agent of mantle metasomatism.
208
Table 4.1: Bulk rock composition of sample B94-26 and of protolith (KMC) andmajor/trace element compositions of major minerals. na — element was notanalysed. bdl — bellow detection limit.
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209
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Chapter 5
Conclusions
In this thesis experiments and the study of natural UHP rocks demonstrates
the important role of monazite in redistribution of REE, Th and U in crustal
rocks. The major results of this work are numerical data such as equation for
monazite solubility, partition coefficients, estimates of melt loss in melting of
metasediments in subduction zone and compositions of UHP melts. These data
can be applied to various petrological problems.
The experimental study of monazite melt partitioning from Chapter 1 pro-
vides a numerical background for the interpretation of REE, Th and U behaviour
in melting processes and thus it is a backbone of the present study. The ex-
periments in the system granite+monazite brought up two types of essential
information: solubility of monazite over a large range of pressures and temper-
atures and monazite/melt partitioning coefficients. Data demonstrated a large
positive temperature effect on the monazite solubility and significant negative
effect of pressure. Comparison with previous experimental data showed that Th
abundance relative to LREE has a significant effect on the LREE solubility and
consistency between different studies can be obtained only if this effect is taken
into account. Compilation of the new data with existing ones resulted in the new
formulation of monazite solubility in granitic melts, which takes into account
temperature, pressure, water content in the melt, and Th content in monazite:
ln∑
LREE = 16.16(±0.3) + 0.23(±0.07)√
H2O
−11494(±410)/T − 19.4(±4)P/T + ln XLREEmnz
Where H2O is in weight percent, T is in Kelvin, P in kbar and∑
LREE is
the sum of La-Sm in ppm; XLREEmnz is the molar ratio of LREE to the sum of all
cations (REE, Th, U) in monazite.
Another primary result of this study is the first experimental data set of mon-
215
216
azite/melt partitioning coefficients for a wide range of trace elements. Monazite
concentrates REE, Y, Th, U, As and V and has the highest preference to LREE
and Th. Rare earth elements heavier than Nd have continuously decreasing parti-
tioning coefficients and U is also significantly less compatible than Th and LREE.
Pressure and temperature increases compatibility of U in monazite but, over all
investigated conditions, U remains significantly less compatible than Th. For
other elements, the effect of pressure and temperature on relative fractionation
by monazite is small. The reduced compatibility of U in monazite has far reach-
ing geochemical implications; it results in overall incompatible behaviour of U
in melting with residual monazite and allows for different melting regimes to be
distinguished. This thesis demonstrates that the new experimental data can be
applied for interpretation of geochemical data of high grade rocks, melt inclusions
and geochronology.
Chapter 2 presents a detailed petrologic and geochronologic investigation of
four metasedimentary samples from the Barchi Kol’ area of the Kokchetav com-
plex. For a first time for a set of samples from the Kokchetav complex trace
element zoning of garnet and accessory minerals were studied together with min-
eral inclusions in zircon and monazite.
The Barchi Kol’ terrain of the Kokchetav complex is mostly composed of
UHP rocks, but also contains rocks with different PT paths, and lower peak con-
ditions, which experienced metamorphism coeval with other UHP rocks. It is
suggeste3d that these rocks represent different portions of the subduction chan-
nel and demonstrate that the UHP terrain is not coherent, but contains slices
or blocks of foreign material. LREE, Th and U in metapelites which did not
underwent UHP melting were hosted in monazite, which recrystallized at vari-
ous stages of the metamorphic evolution. PT paths of the Kokchetav samples
bring evidences of large temperature gradients in HP/UHP rocks which are being
exhumed and record rapid increase of temperature at the depth of 20-30 kbar,
thus supporting the existence of a “could nose” of mantle wedge, as predicted by
geodynamic models.
Metasedimentary UHP rocks from the Kokchetav complex are the perfect
object for a case study of the effect of partial melting on composition of crustal
rocks in subduction zones. The main feature of the Kokchetav gneisses is a strong
depletion in LREE, Th and U (Chapter 3). This feature is so different from
typical crustal melting that, in my opinion, it is worth to defining Kokchetav
type melting as melting of crustal rocks when the host phase for LREE, Th and
U (monazite/allanite) is completely dissolved in the melt and melt is efficiently
217
extracted. The Kokchetav type melting has important geochemical implications:
it produces large fractionation of LREE, Th and U from other incompatible trace
elements, and melting results in the redistribution of heat producing elements
(Th, U and partially K), which does not occur in typical anatexis of continental
crust.
Polyphase inclusions are present in some UHP gneisses and experiments on
high pressure homogenization of inclusions demonstrate that inclusions are for-
mer melts and allowed to determine the major and trace element composition of
inclusions (Chapter 4). Polyphase inclusions provide information complimentary
to bulk rock geochemistry and demonstrate that Kokchetav type melting gen-
erates potassic granitic melts with high LREE, Th and U contents. Elemental
ratios such as Th/U and Th/La are not fractionated in this type of melting.
Common melting models assume either that the melt is removed from the
residue instantly (fractional melting), or that the rock melts to a certain degree
and then all melt is completely extracted (batch melting). Alternatively, melt
is extracted when it reaches certain percentage (dynamic melting, Zhou, 2007).
However, these models are inadequate for the Kokchetav gneisses, because ef-
ficiency of melt extraction varies from very high in ultra-depleted samples to
modest in semi-depleted gneisses. A constant, single value of melt extraction
can not be assigned to all samples, and in fact it remain unclear what factors
controlled efficiency of melt extraction. I have developed a new simple melting
model with variable fraction of melt in the residue. The models allows to obtain
phase abundances in Kokchetav UHP gneisses at peak conditions. Therefore,
it is estimated that Kokchetav gneisses lost 10-40 % of granitic melts and the
restite was composed of Coe, Grt, ±Phe, ± Cpx, Zrn, and ± Rt. 50-98 % of melt
generated in gneisses was extracted from the residue. These estimates open a
new way for the interpretation of element behaviour during UHP anatexis, from
experimental data on solubility and partitioning properties of minerals.
The geochemistry of the Kokchetav UHP gneisses shows significantly different
behaviour of elements than is generally accepted. LREE and Th are often con-
sidered as immobile elements, but in Kokchetav type melting they are the most
incompatible elements due to high temperature of melting, sufficient melting de-
gree and efficient extraction of melt in rocks with low melting degree. In contrast,
LILE, which are typically considered as very mobile and incompatible elements,
in the Kokchetav gneisses were only moderately incompatible, because of pres-
ence of phengite and low efficiency of melt extraction at high degrees of melting.
Nb and Ta were moderately incompatible during melting and were significantly
218
fractionated.
Kokchetav type melting is fundamentally different from the normal crustal
melting where there is no depletion in LREE in residual rocks. Is Kokchetav type
melting common in the deep Earth? Does it occur in the normal subduction zone
setting as recently proposed by Behn et al. (2011); Hacker et al. (2011)? How are
related melting and exhumation of the Kokchetav gneisses? These are important
open questions for further studies.
Bibliography
Acosta-Vigil, A., Buick, I., Hermann, J., Cesare, B., Rubatto, D., London, D. &
Morgan, G. B. (2010). Mechanisms of crustal anatexis: a geochemical study
of partially melted metapelitic enclaves and host dacite, SE Spain. Journal of
Petrology 51, 785–821.
Acosta-Vigil, A., London, D., Morgan, G. & Dewers, T. (2006). Dissolution
of quartz, albite, and orthoclase in H2O-saturated haplogranitic melt at 800
degrees C and 200 MPa: Diffusive transport properties of granitic melts at
crustal anatectic conditions. Journal of Petrology 47, 231–254.
Aleinikoff, J., Schenck, W., Plank, M., Srogi, L., Fanning, C., Kamo, S. & Bos-
byshell, H. (2006). Deciphering igneous and metamorphic events in high-grade
rocks of the Wilmington Complex, Delaware: Morphology, cathodolumines-
cence and backscattered electron zoning, and SHRIMP U-Pb geochronology of
zircon and monazite. Geological Society of America Bulletin 118, 39–64.
Armbruster, T., Bonazzi, P., Akasaka, M., Bermanec, V., Chopin, C., Giere, R.,
Heuss-Assbichler, S., Liebscher, A., Menchetti, S., Pan, Y. & Pasero, M. (2006).
Recommended nomenclature of epidote-group minerals. European Journal of
Mineralogy 18, 551–567.
Auzanneau, E., Schmidt, M. W., Vielzeuf, D. & Connolly, J. A. D. (2010). Tita-
nium in phengite: a geobarometer for high temperature eclogites. Contributions
to Mineralogy and Petrology 159, 1–24.
Auzanneau, E., Vielzeuf, D. & Schmidt, M. W. (2006). Experimental evidence
of decompression melting during exhumation of subducted continental crust.
Contributions to Mineralogy and Petrology 152, 125–148.
Ayers, J. & Watson, E. (1991). Solubility of apatite, monazite, zircon, and
rutile in supercritical aqueous fluids with implications for subduction zone
geochemistry. Philosophical Transactions of the Royal Society of London Series
A-Mathematical Physical and Engineering Sciences 335, 365–375.
219
220
Bali, E., Audetat, A. & Keppler, H. (2011). The mobility of U and Th in sub-
duction zone fluids: an indicator of oxygen fugacity and fluid salinity. Contri-
butions to Mineralogy and Petrology 161, 597–613.
Bartoli, O., Cesare, B., Poli, S., Bodnar, R., Frezzotti, M., Acosta-Vigil, A. &
Meli, S. (2011). Melting in the Deep Crust: Message from Melt Inclusions in
Peritectic Garnet from Migmatites. Mineralogical Magazine 75, 495. Gold-
schmidt conference 2011.
Bau, M. (1996). Controls on the fractionation of isovalent trace elements in
magmatic and aqueous systems: Evidence from Y/Ho, Zr/Hf, and lanthanide
tetrad effect. Contributions to Mineralogy and Petrology 123, 323–333.
Bea, F. (1996). Residence of REE, Y, Th and U in granites and crustal protoliths;
Implications for the chemistry of crustal melts. Journal of Petrology 37, 521–
552.
Bea, F. & Montero, P. (1999). Behavior of accessory phases and redistribu-
tion of Zr, REE, Y, Th, and U during metamorphism and partial melting of
metapelites in the lower crust: An example from the Kinzigite Formation of
Ivrea-Verbano, NW Italy. Geochimica et Cosmochimica Acta 63, 1133–1153.
Bea, F., Pereira, M. & Stroh, A. (1994). Mineral leucosome trace-element parti-
tioning in a peraluminous migmatite (a laser ablation-ICP-MS study). Chem-
ical Geology 117, 291–312.
Beattie, P., Drake, M., Jones, J., Leeman, W., Longhi, J., Mckay, G., Nielsen,
R., Palme, H., Shaw, D., Takahashi, E. & Watson, B. (1993). Terminology for
trace-element partitioning. Geochimica et Cosmochimica Acta 57, 1605–1606.
Bebout, G., Ryan, J. & Leeman, W. (1993). B-be systematics in subduction-
related metamorphic rocks - characterization of the subducted component.
Geochimica Et Cosmochimica Acta 57, 2227–2237.
Bebout, G., Ryan, J., Leeman, W. & Bebout, A. (1999). Fractionation of trace
elements by subduction-zone metamorphism - effect of convergent-margin ther-
mal evolution. Earth And Planetary Science LetterS 171, 63–81.
Bebout, G. E., Bebout, A. E. & Graham, C. M. (2007). Cycling of B, Li, and
LILE (K, Cs, Rb, Ba, Sr) into subduction zones: SIMS evidence from micas
in high-P/T metasedimentary rocks. Chemical Geology 239, 284–304.
221
Becker, H., Jochum, K. & Carlson, R. (2000). Trace element fractionation during
dehydration of eclogites from high-pressure terranes and the implications for
element fluxes in subduction zones. Chemical Geology 163, 65–99.
Behn, M. D., Kelemen, P. B., Hirth, G., Hacker, B. R. & Massonne, H.-J. (2011).
Diapirs as the source of the sediment signature in arc lavas. Nature Geoscience
4, 641–646.
Beltrando, M., Compagnoni, R. & Lombardo, B. (2010). (Ultra-) High-pressure
metamorphism and orogenesis: An Alpine perspective. Gondwana Research
18, 147–166.
Bingen, B., Demaiffe, D. & Hertogen, J. (1996). Redistribution of rare earth
elements, thorium, and uranium over accessory minerals in the course of am-
phibolite to granulite facies metamorphism: The role of apatite and monazite
in orthogneisses from southwestern Norway. Geochimica et Cosmochimica Acta
60, 1341–1354.
Black, L., Kamo, S., Allen, C., Aleinikoff, J., Davis, D., Korsch, R. & Foudoulis,
C. (2003). TEMORA 1: a new zircon standard for Phanerozoic U-Pb
geochronology. Chemical Geology 200, 155–170.
Blundy, J., Robinson, J. & Wood, B. (1998). Heavy ree are compatible in clinopy-
roxene on the spinel lherzolite solidus. Earth and Planetary Science Letters
160, 493–504.
Blundy, J. & Wood, B. (1994). Prediction of crystal-melt partition-coefficients
from elastic-moduli. Nature 372, 452–454.
Buslov, M. M., Ryabinin, A. B., Zhimulev, F. I. & Travin, A. V. (2009). Man-
ifestations of the Late Carboniferous and Early Permian stages of formation
of nappe-fold structures in the southern framework of the Siberian platform
(East Sayany, South Siberia). Doklady Earth Sciences 428, 1105–1108.
Buslov, M. M. & Vovna, G. M. (2008). Composition and geodynamic nature
of the protoliths of diamondiferous rocks from the Kumdy-Kol deposit of the
Kokchetav metamorphic belt, northern Kazakhstan. Geochemistry Interna-
tional 46, 887–896.
Buslov, M. M., Zhimulev, F. I. & Travin, A. V. (2010). New data on the struc-
tural setting and Ar-40/Ar-39 age of the MP-LP metamorphism of the Daulet
222
formation, Kokchetav metamorphic belt, Northern Kazakhstan, and their tec-
tonic interpretation. Doklady Earth Sciences 434, 1147–1151.
Carswell, D., Cuthbert, S. & Ravna, E. (1999). Ultrahigh-pressure metamorphism
in the Western Gneiss Region of the Norwegian Caledonides. International
Geology Review 41, 955–966.
Cartigny, P., De Corte, K., Shatsky, V., Ader, M., De Paepe, P., Sobolev, N. &
Javoy, M. (2001). The origin and formation of metamorphic microdiamonds
from the Kokchetav massif, Kazakhstan: a nitrogen and carbon isotopic study.
Chemical Geology 176, 265–281.
Castro, A., Garcia-Casco, A., Fernandez, C., Corretge, L. G., Moreno-Ventas, I.,
Gerya, T. & Loew, I. (2009). Ordovician ferrosilicic magmas: Experimental
evidence for ultrahigh temperatures affecting a metagreywacke source. Gond-
wana Research 16, 622–632.
Cesare, B., Ferrero, S., Salviolo-Mariani, E., Pedron, D. & Cavallo, A. (2009).
“Nanogranite” and glassy inclusions: The anatectic melt in migmatites and
granulites. Geology 37, 627–630.
Chauvel, C., Lewin, E., Carpentier, M., Arndt N.T. & Marini J.-C. (2008). Role
of recycled oceanic basalt and sediment in generating the Hf-Nd mantle array.
Nature Geoscience pages 64–67.
Cherniak, D. J. & Watson, E. B. (2007). Ti diffusion in zircon. Chemical Geology
242, 473–483.
Chopin, C. (1987). Very-high-pressure metamorphism in the Western Alps -
implications for subduction of continental-crust. Philosophical Transactions Of
The Royal Society Of London Series A-Mathematical Physical And Engineering
Sciences 321, 183–197.
Chopin, C. (2003). Ultrahigh-pressure metamorphism: tracing continental crust
into the mantle. Earth and Planetary Science Letters 212, 1–14.
Chopin, C., Brunet, F., Gebert, W., Medenbach, O. & Tillmanns, E. (1993).
Bearthite, Ca2Al[PO4]2(OH), a new mineral from high-pressure terranes of
the Western Alps. Schweizerische Mineralogische Und Petrographische Mit-
teilungen 73, 1–9.
223
Claoue-Long, J., Sobolev, N., Shatsky, V. & Sobolev, A. (1991). Zircon response
to diamond-pressure metamorphism in the Kokchetav massif, USSR. Geology
19, 710–713.
Cloos, M. (1982). Flow melanges - numerical modeling and geologic constraints
on their origin in the Franciscan subduction complex, California. Geological
Society Of America Bulletin 93, 330–345.
Condie, K. (1993). Chemical-composition and evolution of the upper continental-
crust - contrasting results from surface samples and shales. Chemical Geology
104, 1–37.
Connolly, J. (2005). Computation of phase equilibria by linear programming: a
tool for geodynamic modeling and its application to subduction zone decar-
bonation. Earth and Planetary Science Letters 236, 524–541.
Corfu, F., Hanchar, J., Hoskin, P. & Kinny, P. (2003). Atlas of zircon textures.
In: Hanchar, JM and Hoskin, PWO (ed.) Zircon, volume 53 of Reviews In
Mineralogy & Geochemistry, pages 469–500. Mineral Soc Amer, Washington,
DC, USA: Mineralogical Soc America.
Cox, R., Lowe, D. & Cullers, R. (1995). The influence of sediment recycling and
basement composition on evolution of mudrock chemistry in the southwestern
united-states. Geochimica et Cosmochimica Acta 59, 2919–2940.
Cressey, G., Wall, F. & Cressey, B. A. (1999). Differential REE uptake by sector
growth of monazite. Mineralogical Magazine 63, 813 – 828.
Crowley, J. L., Brown, R. L., Gervais, F. & Gibson, H. D. (2008). Assessing Inher-
itance of Zircon and Monazite in Granitic Rocks from the Monashee Complex,
Canadian Cordillera. Journal Of Petrology 49, 1915–1929.
Cuney, M. & Friedrich, M. (1987). Physicochemical and crystal-chemical con-
trols on accessory mineral paragenesis in granitoids - implications for uranium
metallogenesis. Bulletin De Mineralogie 110, 235–247.
Cuthbert, S., Carswell, D., Krogh-Ravna, E. & Wain, A. (2000). Eclogites and
eclogites in the Western Gneiss Region, Norwegian Caledonides. Lithos 52,
165–195.
224
De Corte, K., Korsakov, A., Taylor, W. R., Cartigny, P., Ader, M. & De Paepe,
P. (2000). Diamond growth during ultrahigh-pressure metamorphism of the
Kokchetav Massif, northern Kazakhstan. Island Arc 9, 428 – 438.
Devine, J., Gardner, J., Brack, H., Layne, G. & Rutherford, M. (1995). Compar-
ison of microanalytical methods for estimating H2O contents of silicic volcanic
glasses. American Mineralogist 80, 319–328.
Dobretsov, N., Buslov, M. & Zhimulev, F. (2005). Cambrian-Ordovician tectonic
evolution of the Kokchetav metamorphic belt, northern Kazakhstan. Russian
Geology and Geophysics 46, 785–795.
Dobretsov, N., Buslov, M., Zhimulev, F., Travin, A. & Zayachkovsky, A. (2006).
Vendian-Early Ordovician geodynamic evolution and model for exhumation of
ultrahighand high-pressure rocks from the Kokchetav subduction-collision zone
(northern Kazakhstan). Russian Geology and Geophysics 47, 424–440.
Dobretsov, N. & Shatsky, V. (2004). Exhumation of high-pressure rocks of the
Kokchetav massif: facts and models. Lithos 78, 307–318.
Dobretsov, N., Sobolev, N., Shatsky, V., Coleman, R. & Ernst, W. (1995). Geo-
tectonic evolution of diamondiferous paragneisses of the Kokchetav Complex,
Northern Kazakhstan - the geologic enigma of ultrahigh-pressure crustal rocks
within Phanerozoic foldbelt. The Island Arc 4, 267–279.
Dobretsov, N., Theunissen, K., Dobretsov, N. & Smirnova, L. (1999). Field Guide
Book of the IV International Eclogite Field Symposium, chapter Geological and
tectonic outline of the Kokchetav massif, pages 6–24. Novosibirsk, Russia.
Dobretsov, N., Theunissen, K. & Smirnova, L. (1998). Structural and geodynamicevolution of the diamondiferous metamorphic rocks of the Kokchetav massif(Kazakhstan). Geologiya I Geofizika 39, 1645–1666.
Dobrzhinetskaya, L., Braun, T., Sheshkel, G. & Podkuiko, Y. (1994). Geology
and structure of diamond-bearing rocks of the Kokchetav massif (Kazakhstan).
Tectonophysics 233, 293–313.
Douce, A. & Beard, J. (1995). Dehydration-melting of biotite gneiss and quartz
amphibolite from 3 to 15 kbar. Journal of Petrology 36, 707–738.
225
Ducea, M., Lutkov, V., Minaev, V., Hacker, B., Ratschbacher, L., Luffi, P.,
Schwab, M., Gehrels, G., McWilliams, M., Vervoort, J. & Metcalf, J. (2003).
Building the Pamirs: The view from the underside. Geology 31, 849–852.
Eggins, S., Rudnick, R. & McDonough, W. (1998). The composition of peridotites
and their minerals: A laser-ablation ICP-MS study. Earth and Planetary
Science Letters 154, 53–71.
Engvik, A., Austrheim, H. & Erambert, M. (2001). Interaction between fluid
flow, fracturing and mineral growth during eclogitization, an example from the
Sunnfjord area, Western Gneiss Region, Norway. Lithos 57, 111–141.
Ferrando, S., Frezzotti, M. L., Petrelli, M. & Compagnoni, R. (2009). Metasoma-
tism of continental crust during subduction: the UHP whiteschists from the
Southern Dora-Maira Massif (Italian Western Alps). Journal of Metamorphic
Geology 27, 739–756.
Ferry, J. M. & Watson, E. B. (2007). New thermodynamic models and revised
calibrations for the Ti-in-zircon and Zr-in-rutile thermometers. Contributions
to Mineralogy and Petrology 154, 429–437.
Finger, F. & Krenn, E. (2007). Three metamorphic monazite generations in a
high-pressure rock from the Bohemian Massif and the potentially important
role of apatite in stimulating polyphase monazite growth along a PT loop.
Lithos 95, 103–115.
Foley, S., Barth, M. & Jenner, G. (2000). Rutile/melt partition coefficients for
trace elements and an assessment of the influence of rutile on the trace element
characteristics of subduction zone magmas. Geochimica et Cosmochimica Acta
64, 933–938.
Forster, H. (1998). The chemical composition of REE-Y-Th-U-rich accessory
minerals in peraluminous granites of the Erzgebirge-Fichtelgebirge region, Ger-
many, Part I: The monazite-(Ce)-brabantite solid solution series. American
Mineralogist 83, 259–272.
Forster, H. (2000). Cerite-(Ce) and thorian synchysite-(Ce) from the Nieder-
bobritzsch granite, Erzgebirge, Germany: Implications for the differential mo-
bility of the LREE and Th during alteration. Canadian Mineralogist 38, 67–79.
226
Forster, M., Lister, G., Compagnoni, R., Giles, D., Hills, Q., Betts, P., Bel-
trando, M., Tamagno, E. & In: Pasquare G. and Venturini C. (Eds), . (2004).
Mapping of oceanic crust with ”HP” to ”UHP” metamorphism: The Lago di
Cignana Unit, (Western Alps), chapter Mapping Geology in Italy, pages 279–
288. Firenze: SELCA.
Frei, D., Liebscher, A., Wittenberg, A. & Shaw, C. (2003). Crystal chemical
controls on rare earth element partitioning between epidote-group minerals and
melts: an experimental and theoretical study. Contributions to Mineralogy and
Petrology 146, 192–204.
Fulmer, E. C., Nebel, O. & van Westrenen, W. (2010). High-precision high field
strength element partitioning between garnet, amphibole and alkaline melt
from Kakanui, New Zealand. Geochimica et Cosmochimica Acta 74, 2741–
2759.
Gao, X.-Y., Zheng, Y.-F. & Chen, Y.-X. (2011). U-Pb ages and trace elements
in metamorphic zircon and titanite from UHP eclogite in the Dabie orogen:
constraints on P-T-t path. Journal of Metamorphic Geology 29, 721–740.
Gerya, T. (2011). Future directions in subduction modeling. Journal Of Geody-
namics 52, 344–378.
Giere, R. & Sorensen, S. (2004). Allanite and other REE-rich epidote-group
minerals. In: Epidotes, volume 56 of Reviews In Mineralogy & Geochemistry,
pages 431–493. Washington.
Goldschmidt, V. & Peters, C. (1933). Uber die Anreicherung seltener Element
in Steinkohlen. Nachr. Ges. Wiss. Gottingen 4, 371–386. (in German).
Gordon, S., Kelemen, P., Hacker, B., Luffi, P. & Ratschbacher, L. (2011). Partial
Melting and its Role in Elemental Recycling: Insight from Pamir Metasedi-
mentary Xenoliths. In: Goldschmidt Conference 2011 Abstracts, page 937.
Gratz, R. & Heinrich, W. (1997). Monazite-xenotime thermobarometry: Experi-
mental calibration of the miscibility gap in the binary system CePO4−Y PO4.
American Mineralogist 82, 772–780.
Green, T. & Adam, J. (2002). Pressure effect on Ti- or P-rich accessory mineral
saturation in evolved granitic melts with differing K2O/Na2O ratios. Lithos
61, 271–282.
227
Green, T. & Hellman, P. (1982). Fe-Mg partitioning between coexisting garnet
and phengite at high-pressure, and comments on a garnet phengite geother-
mometer. Lithos 15, 253–266.
Gutscher, M., Maury, R., Eissen, J. & Bourdon, E. (2000). Can slab melting be
caused by flat subduction? Geology 28, 535–538.
Hacker, B., Luffi, P., Lutkov, V., Minaev, V., Ratschbacher, L., Plank, T., Ducea,
M., Patino-Douce, A., McWilliams, M. & Metcalf, J. (2005). Near-ultrahigh
pressure processing of continental crust: Miocene crustal xenoliths from the
Pamir. Journal of Petrology 46, 1661–1687.
Hacker, B. R., Andersen, T. B., Johnston, S., Kylander-Clark, A. R. C., Peter-
man, E. M., Walsh, E. O. & Young, D. (2010). High-temperature deformation
during continental-margin subduction & exhumation: The ultrahigh-pressure
Western Gneiss Region of Norway. Tectonophysics 480, 149–171.
Hacker, B. R., Kelemen, P. B. & Behn, M. D. (2011). Differentiation of the
continental crust by relamination. Earth and Planetary Science Letters 307,
501–516.
Halter, W., Pettke, T., Heinrich, C. & Rothen-Rutishauser, B. (2002). Major to
trace element analysis of melt inclusions by laser-ablation ICP-MS: methods
of quantification. Chemical Geology 183, 63–86.
Hanson, G. & Langmuir, C. (1978). Modelling of major elements in mantle-melt
systems using trace element aproaches. Geochimica et Cosmochimica Acta 42,
725–741.
Harlov, D. E., Wirth, R. & Hetherington, C. J. (2007). The relative stability
of monazite and huttonite at 300-900 oC and 200-1000 mpa: metasomatism
and the propagation of metastable mineral phases. American Mineralogist 92,
1652–1664.
Hawkesworth, C., Hergt, J., Ellam, R. & Mcdermott, F. (1991). Element fluxes
associated with subduction related magmatism. Philosophical Transactions of
the Royal Society of London Series A-Mathematical Physical and Engineering
Sciences 335, 393–405.
Helsel, D. & Hirsch, R. (2002). Statistical methods in water resources, volume
Book 4, Chapter A3 of Techniques of Water-Resources Investigations. U.S.
Geological Survey.
228
Hermann, J. (2002a). Allanite: thorium and light rare earth element carrier in
subducted crust. Chemical Geology 192, 289 – 306.
Hermann, J. (2002b). Experimental constraints on phase relations in subducted
continental crust. Contributions To Mineralogy And Petrology 143, 219 – 235.
Hermann, J. (2003). Experimental evidence for diamond-facies metamorphism
in the Dora-Maira massif. Lithos 70, 163 – 182.
Hermann, J. & Green, D. H. (2001). Experimental constraints on high pressure
melting in subducted crust. Earth And Planetary Science Letters 188, 149 –
168.
Hermann, J. & Rubatto, D. (2003). Relating zircon and monazite domains to
garnet growth zones: age and duration of granulite facies metamorphism in
the Val Malenco lower crust. Journal Of Metamorphic Geology 21, 833 – 852.
Hermann, J. & Rubatto, D. (2009). Accessory phase control on the trace element
signature of sediment melts in subduction zones. Chemical Geology 265, 512
– 526.
Hermann, J. & Rubatto, D. (2012). Subduction of continental crust to mantle
depth: Geochemistry of ultrahigh-pressure rocks. In: Treatise on Geochem-
istry. Elsevier.
Hermann, J., Rubatto, D., Korsakov, A. & Shatsky, V. S. (2001). Multiple zircon
growth during fast exhumation of diamondiferous, deeply subducted continen-
tal crust (Kokchetav Massif, Kazakhstan). Contributions To Mineralogy And
Petrology 141, 66 – 82.
Hermann, J., Spandler, C., Hack, A. & Korsakov, A. V. (2006). Aqueous fluids
and hydrous melts in high-pressure and ultra-high pressure rocks: Implications
for element transfer in subduction zones. Lithos 92, 399 – 417.
Hermann, J. & Spandler, C. J. (2008). Sediment melts at sub-arc depths: An
experimental study. Journal Of Petrology 49, 717 – 740.
Hodges, K. & Crowley, P. (1985). Error estimation and empirical geothermo-
barometry for pelitic systems. American Mineralogist 70, 702–709.
Hoffman, D., Lawrence, F., Mewherte.JL & Rourke, F. (1971). Detection of
plutonium-244 in nature. Nature 234, 132–&.
229
Hofmann, A. (1988). Chemical differentiation of the Earth - the relationship
between mantle, continental-crust, and oceanic-crust. Earth and Planetary
Science Letters 90, 297–314.
Hogan, J. & Sinha, A. (1991). The effect of accessory minerals on the redis-
tribution of lead isotopes during crustal anatexis: a model. Geochimica et
Cosmochimica Acta 55, 335–348.
Holland, T. & Powell, R. (1998). An internally consistent thermodynamic data
set for phases of petrological interes. Journal of Metamorphic Geology 16,
309–343.
Holloway, J. & Wood, B. (1988). Simulating the Earth - Experimental geochem-
istry. Springer. 288 p.
Honda, S., Gerya, T. & Zhu, G. (2010). A simple three-dimensional model of
thermo-chemical convection in the mantle wedge. Earth and Planetary Science
Letters 290, 311–318.
Huang, W. & Wyllie, P. (1981). Phase-relationships of S-type granite with H2O
to 35 kbar - muscovite granite from Harney Peak, South-Dakota. Journal Of
Geophysical Research 86, 515–529.
Hwang, S., Chu, H., Yui, T., Shen, P., Schertl, H., Lion, J. & Sobolev, N. (2006).
Nanometer-size P/K-rich silica glass (former melt) inclusions in microdiamond
from the gneisses of Kokchetav and Erzgebirge massifs: Diversified character-
istics of the formation media of metamorphic microdiamond in UHP rocks due
to host-rock buffering. Earth and Planetary Science Letters 243, 94–106.
Hwang, S., Shen, P., Chu, H., Yui, T. & Lin, C. (2001). Genesis of microdiamonds
from melt and associated multiphase inclusions in garnet of ultrahigh-pressure
gneiss from Erzgebirge, Germany. Earth and Planetary Science Letters 188,
9–15.
Hwang, S., Shen, P., Chu, H., Yui, T., Liou, J., Sobolev, N. & Shatsky, V.
(2005). Crust-derived potassic fluid in metamorphic microdiamond. Earth and
Planetary Science Letters 231, 295–306.
Hwang, S., Shen, P., Chu, H., Yui, T., Liou, J., Sobolev, N., Zhang, R., Shatsky,
V. & Zayachkovsky, A. (2004). Kokchetavite: a new potassium-feldspar poly-
morph from the Kokchetav ultrahigh-pressure terrane. Contributions to Min-
eralogy and Petrology 148, 380–389.
230
Hwang, S., Shen, P., Yui, T. & Chu, H. (2003). Metal-sulfur-COH-silicate fluid
mediated diamond nucleation in Kokchetav ultrahigh-pressure gneiss. Euro-
pean Journal of Mineralogy 15, 503–511.
Hwang, S.-L., Shen, P., Chu, H.-T., Yui, T.-F., Liou, J. G. & Sobolev,
N. V. (2009). Kumdykolite, an orthorhombic polymorph of albite, from the
Kokchetav ultrahigh-pressure massif, Kazakhstan. European Journal of Min-
eralogy 21, 1325–1334.
Ionov, D. & Harmer, R. E. (2002). Trace element distribution in calcite-dolomite
carbonatites from Spitskop: inferences for differentiation of carbonatite mag-
mas and the origin of carbonates in mantle xenoliths. Earth and Planetary
Science Letters 198, 495 – 510.
Irber, W. (1999). The lanthanide tetrad effect and its correlation with K/Rb,
Eu/Eu*, Sr/Eu, Y/Ho, and Zr/Hf of evolving peraluminous granite suites.
Geochimica et Cosmochimica Acta 63, 489–508.
Ishikawa, M., Kaneko, Y., Anma, R. & Yamamoto, H. (2000). Subhorizontal
boundary between ultrahigh-pressure and low-pressure metamorphic units in
the Sulu-Tjube area of the Kokchetav Massif, Kazakhstan. Island Arc 9, 317
– 327.
Janots, E., Negro, F., Brunet, F., Goffe, B., Engi, M. & Bouybaouene, M. (2006).
Evolution of the REE mineralogy in HP-LT metapelites of the Sebtide complex,
Rif, Morocco: Monazite stability and geochronology. Lithos 87, 214–234.
Janousek, V., Finger, F., Roberts, M., Fryda, J., Pin, C. & Dolejs, D. (2004).
Deciphering the petrogenesis of deeply buried granites: whole-rock geochem-
ical constraints on the origin of largely undepleted felsic granulites from the
Moldanubian Zone of the Bohemian Massif. Transactions Of The Royal Society
Of Edinburgh-Earth Sciences 95, 141–159.
Jochum, K., Stolz, A. & McOrist, G. (2000). Niobium and tantalum in car-
bonaceous chondrites: Constraints on the solar system and primitive mantle
niobium/tantalum, zirconium/niobium, and niobium/uranium ratios. Mete-
oritics & Planetary Science 35, 229–235.
John, T. & Schenk, V. (2003). Partial eclogitisation of gabbroic rocks in a late
Precambrian subduction zone (Zambia): prograde metamorphism triggered by
fluid infiltration. Contributions to Mineralogy and Petrology 146, 174–191.
231
John, T., Scherer, E., Haase, K. & Schenk, V. (2004). Trace element fractionation
during fluid-induced eclogitization in a subducting slab: trace element and
Lu-Hf-Sm-Nd isotope systematics. Earth and Planetary Science Letters 227,
441–456.
Jones, A., Wall, F. & Williams, C. (1996). Rare earth minerals: chemistry, origin
and ore deposits. The Mineralogical Society Series, Vol. 7.
Kaneko, Y., Maruyama, S., Terabayashi, M., Yamamoto, H., Ishikawa, M.,
Anma, R., Parkinson, C. D., Ota, T., Nakajima, Y., Katayama, I., Yamamoto,
J. & Yamauchi, K. (2000). Geology of the Kokchetav UHP-HP metamorphic
belt, Northern Kazakhstan. Island Arc 9, 264 – 283.
Katayama, I. & Maruyama, S. (2009). Inclusion study in zircon from ultrahigh-
pressure metamorphic rocks in the Kokchetav massif: an excellent tracer of
metamorphic history. Journal of the Geological Society 166, 783–796.
Katayama, I., Maruyama, S., Parkinson, C., Terada, K. & Sano, Y. (2001).
Ion micro-probe U-Pb zircon geochronology of peak and retrograde stages of
ultrahigh-pressure metamorphic rocks from the Kokchetav massif, northern
Kazakhstan. Earth and Planetary Science Letters 188, 185–198.
Kelley, K., Plank, T., Farr, L., Ludden, J. & Staudigel, H. (2005). Subduction
cycling of U, Th, and Pb. Earth and Planetary Science Letters 234, 369–383.
Kelsey, D. E., Clark, C. & Hand, M. (2008). Thermobarometric modelling of
zircon and monazite growth in melt-bearing systems: examples using model
metapelitic and metapsammitic granulites. Journal of Metamorphic Geology
26, 199–212.
Kennedy, M., Droser, M., Mayer, L., Pevear, D. & Mrofka, D. (2006). Late
Precambrian oxygenation; Inception of the clay mineral factory. Science 311,
1446–1449.
Keppler, H. & Wyllie, P. (1990). Role of fluids in transport and fractionation of
uranium and thorium in magmatic processes. Nature 348, 531–533.
Klimm, K., Blundy, J. D. & Green, T. H. (2008). Trace element partitioning
and accessory phase saturation during H2O-saturated melting of basalt with
implications for subduction zone chemical fluxes. Journal of Petrology 49,
523–553.
232
Korsakov, A. V. & Hermann, J. (2006). Silicate and carbonate melt inclusions
associated with diamonds in deeply subducted carbonate rocks. Earth And
Planetary Science Letters 241, 104 – 118.
Korsakov, A. V., Perraki, M., Zhukov, V. P., De Gussem, K., Vandenabeele,
P. & Tomilenko, A. A. (2009). Is quartz a potential indicator of ultrahigh-
pressure metamorphism? Laser Raman spectroscopy of quartz inclusions in
ultrahigh-pressure garnets. European Journal of Mineralogy 21, 1313–1323.
Korsakov, A. V., Shatsky, V. S., Sobolev, N. V. & Zayachokovsky, A. A. (2002).
Garnet-biotite-clinozoisite gneiss: a new type of diamondiferous metamorphic
rock from the Kokchetav Massif. European Journal of Mineralogy 14, 915 –
928.
Korsakov, A. V., Theunissen, K., Kozmenko, O. A. & Ovchinnikov, Y. I. (2006).
Reaction textures in clinozoisite gneisses. Russian Geology and Geophysics 47,
497 – 510.
Korsakov, A. V., Theunissen, K. & Smirnova, L. V. (2004). Intergranular dia-
monds derived from partial melting of crustal rocks at ultrahigh-pressure meta-
morphic conditions. Terra Nova 16, 146 – 151.
Kretz, R. (1983). Symbols for rock-forming minerals. American Mineralogist 68,
277–279.
Kushev, V. G. & Vinogradov, D. P. (1978). Metamorphic Eclogites. Nauka,
Novosibirsk.
Lang, H. M. & Gilotti, J. A. (2007). Partial melting of metapelites at ultrahigh-
pressure conditions, Greenland Caledonides. Journal of Metamorphic Geology
25, 129–147.
Lavrova, L. D., Pechnikov, V. A., Petrova, M. & Zayachkovsky, A. (1996). Geol-
ogy of the Barchi diamondiferous area. Otechestvennaya Geologiya 12, 12–27.
(in Russian).
Leeman, W. & Phelps, D. (1981). Partitioning of Rare-Earths and other trace-
elements between sanidine and coexisting volcanic glass. Journal Of Geophys-
ical Research 86, 193–199.
233
Letnikov, F., Kostitsyn, Y., Vladykin, N., Zayachkovskii, A. & Mishina, E.
(2004). Isotopic characteristics of the Krasnyi Mai ultramafic alkaline rock
complex, northern Kazakhstan. Doklady Earth Sciences 399A, 1315–1319.
Letnikov, F. A. (2008). Topaz granites in northern Kazakhstan. Petrology 16,
319–334.
Letnikov, F. A., Kotov, A. B., Sal’nikova, E. B., Shershakova, M. M., Shershakov,
A. V., Rizvanova, N. G. & Makeev, A. F. (2007). Granodiorites of the Grenville
phase in the Kokchetav Block, northern Kazakhstan. Doklady Earth Sciences
417, 1195–1197.
Lexa, O., Schulmann, K., Janousek, V., Stipska, P., Guy, A. & Racek, M.
(2011). Heat sources and trigger mechanisms of exhumation of HP granulites
in Variscan orogenic root. Journal Of Metamorphic Geology 29, 79–102.
Liou, J., Zhang, R., Katayama, I., Maruyama, S. & Ernst, W. (2002). Petro-
tectonic characterization f the Kokchetav Massif and the Dabie-Sulu terranes
Ultrahigh-P metamorphism in the so-called P-T Forbidden-Zone. Western
Pacific Earth Sciences 2, 119–148.
Liu, F., Xu, Z., Liou, J., Dong, H. & Xue, H. (2007). Ultrahigh-pressure mineral
assemblages in zircons from surface to 5158 m depth cores in the main drill
hole of Chinese Continental Scientific Drilling Project, southwestern Sulu belt,
China. International Geological Review 49, 454–478.
Liu, F., Xu, Z., Liou, J., Katayama, I., Masago, H., Maruyama, S. & Yang,
J. (2002). Ultrahigh-pressure mineral inclusions in zircons from gneissic core
samples of the Chinese Continental Scientific Drilling Site in eastern China.
European Journal of Mineralogy 14, 499–512.
Liu, F., Xua, Z., Katayama, I., Yang, J., Maruyama, S. & Liou, J. (2001). Mineral
inclusions in zircons of para- and orthogneiss from pre-pilot drillhole CCSD-
PP1, Chinese Continental Scientific Drilling Project. Lithos 59, 199–215.
Ludwig, K. (2003). User’s manual for Isoplot 3.00. A geochronological Toolkit for
Microsoft Excel. Number No. 4a in Special Publication. Berkeley, California:
Berkeley Geochronology Center. Pp. 71.
Lyakhovich, V. & Barinskii, R. (1961). Characteristics of the rare-earth assem-
blages in the accessory minerals of granitoids. Geochemistry 6, 495–509.
234
Madyukov, I., Chupin, V. & Kuzmin, D. (2011). Genesis of scapolite from gran-
ulites (lower-crustal xenoliths from the pamir diatremes): results of study of
melt inclusions. Russian Geology and Geophysics 52, 1319 – 1333.
Malaspina, N., Hermann, J. & Scambelluri, M. (2009). Fluid/mineral interaction
in UHP garnet peridotite. Lithos 107, 38 – 52.
Malaspina, N., Hermann, J., Scambelluri, M. & Compagnoni, R. (2006a). Mul-
tistage metasomatism in ultrahigh-pressure mafic rocks from the North Dabie
Complex (China). Lithos 90, 19 – 42.
Malaspina, N., Hermann, J., Scambelluri, M. & Compagnoni, R. (2006b).
Polyphase inclusions in garnet-orthopyroxenite (Dabie Shan, China) as moni-
tors for metasomatism and fluid-related trace element transfer in subduction
zone peridotite. Earth And Planetary Science Letters 249, 173 – 187.
Marschall, H. R., Altherr, R. & Ruepke, L. (2007). Squeezing out the slab -
modelling the release of Li, Be and B during progressive high-pressure meta-
morphism. Chemical Geology 239, 323–335.
Maruyama, S. & Parkinson, C. D. (2000). Overview of the geology, petrology
and tectonic framework of the high-pressure-ultrahigh-pressure metamorphic
belt of the Kokchetav Massif, Kazakhstan. Island Arc 9, 439 – 455.
Masago, H. (2000). Metamorphic petrology of the Barchi-Kol metabasites, west-
ern Kokchetav ultrahigh-pressure-high-pressure massif, northern Kazakhstan.
Island Arc 9, 358–378.
Massonne, H. (1999). A new occurrence of microdiamonds in Quartzofeldspathic
rocks of the Saxonian Erzgebirge, Germany, and their metamorphic Evolution.
In: J. Gurney, J. Gurney, M. Pascoe & S. Richardson (eds.) Proceed. 7th Int.
Kimberlite Conf., volume 2, pages 533–539.
Massonne, H. (2003). A comparison of the evolution of diamondiferous quartz-
rich rocks from the Saxonian Erzgebirge and the Kokchetav Massif: are so-
called diamondiferous gneisses magmatic rocks? Earth and Planetary Science
Letters 216, 347–364.
Mazzone, P. & Haggerty, S. (1989). Peraluminous xenoliths in kimberlite - meta-
morphosed restites produced by partial melting of pelites. Geochimica et Cos-
mochimica Acta 53, 1551–1561.
235
McDonough, W. & Sun, S. (1995). The composition of the earth. Chemical
Geology 120, 223–253.
Melzer, S. & Wunder, B. (2000). Island-arc basalt alkali ratios: Constraints from
phengite-fluid partitioning experiments. Geology 28, 583–586.
Montel, J. (1986). Experimental-determination of the solubility of Ce-monazite
in SiO2−Al2O3−K2O−Na2O melts at 800 oC, 2 kbar, under H2O-saturated
conditions. Geology 14, 659–662.
Montel, J. (1993). A model for monazite/melt equilibrium and application to the
generation of granitic magmas. Chemical Geology 110, 127–146.
Montel, J. & Vielzeuf, D. (1997). Partial melting of metagreywackes .2. Compo-
sitions of minerals and melts. Contributions to Mineralogy and Petrology 128,
176–196.
Munker, C., Pfander, J., Weyer, S., Buchl, A., Kleine, T. & Mezger, K. (2003).
Evolution of planetary cores and the earth-moon system from Nb/Ta system-
atics. Science 301, 84–87.
Nadolinny, V. A., Shatsky, V. S., Kozmenko, O. A., Stepanov, A. S., Palyanov,
Y. N. & Kupriyanov, I. N. (2006). Study of local concentration of single
substitutional nitrogen atoms in microdiamonds from the Kokchetav massif.
European Journal of Mineralogy 18, 739–743.
Naemura, K., Yokoyama, K., Hirajima, T. & Svojtka, M. (2008). Age determi-
nation of thorianite in phlogopite-bearing spinel-garnet peridotite in the Gfohl
Unit, Moldanubian Zone of the Bohemian Massif. Journal Of Mineralogical
And Petrological Sciences 103, 285–290.
Nebel, O., van Westrenen, W., Vroon, P. Z., Wille, M. & Raith, M. M. (2010).
Deep mantle storage of the Earth’s missing niobium in late-stage residual melts
from a magma ocean. Geochimica et Cosmochimica Acta 74, 4392–4404.
Ni, Y., Hughes, J. & Mariano, A. (1995). Crystal-chemistry of the monazite and
xenotime structures. American Mineralogist 80, 21–26.
Norman, M., Griffin, W., Pearson, N., Garcia, M. & O’Reilly, S. (1998). Quanti-
tative analysis of trace element abundances in glasses and minerals: a compar-
ison of laser ablation inductively coupled plasma mass spectrometry, solution
236
inductively coupled plasma mass spectrometry, proton microprobe and electron
microprobe data. Journal Of Analytical Atomic Spectrometry 13, 477–482.
Nozkhin, A. D. & Turkina, O. M. (1993). Geochemistry of granulites of the Kansk
and Saryzhalgaisk complexes. UIGGM SO RAN: Russian Academy of Sciences
Press. (in Russian).
Ogasawara, Y., Fukasawa, K. & Maruyama, S. (2002). Coesite exsolution from
supersilicic titanite in UHP marble from the Kokchetav Massif, northern Kaza-
khstan. American Mineralogist 87, 454–461.
Ondrejka, M., Uher, P., Prsek, J. & Ozdin, D. (2007). Arsenian monazite-(Ce)
and xenotime-(Y), REE arsenates and carbonates from the Tisovec-Rejkovo
rhyolite, Western Carpathians, Slovakia: Composition and substitutions in the
(REE,Y)XO4 system (X = P, As, Si, Nb, S). Lithos 95, 116–129.
O’Neill, H. S. C., Berry, A. J. & Eggins, S. M. (2008). The solubility and oxidation
state of tungsten in silicate melts: Implications for the comparative chemistry
of W and Mo in planetary differentiation processes. Chemical Geology 255,
346–359.
Overstreet, W. (1967). The geologic occurence of monazite. USGS Professional
Paper 530. Washington: U.S. Dept. of the Interior.
Parkinson, C. (2000). Coesite inclusions and prograde compositional zonation
of garnet in whiteschist of the HP-UHPM Kokchetav massif, Kazakhstan: a
record of progressive UHP metamorphism. Lithos 52, 215–233.
Peacock, S., Rushmer, T. & Thompson, A. (1994). Partial melting of subducting
oceanic-crust. Earth And Planetary Science Letters 121, 227–244.
Pearce, N., Perkins, W., Westgate, J., Gorton, M., Jackson, S., Neal, C. & Ch-
enery, S. (1997). A compilation of new and published major and trace element
data for NIST SRM 610 and NIST SRM 612 glass reference materials. Geostan-
dards Newsletter-the Journal of Geostandards and Geoanalysis 21, 115–144.
Pechnikov, V. A. & Kaminsky, F. V. (2008). Diamond potential of metamorphic
rocks in the Kokchetav Massif, northern Kazakhstan. European Journal of
Mineralogy 20, 395–413.
237
Perchuk, A. L., Burchard, M., Maresch, W. V. & Schertl, H. P. (2005). Fluid-
mediated modification of garnet interiors under ultrahigh-pressure conditions.
Terra Nova 17, 545–553.
Perchuk, A. L., Burchard, M., Maresch, W. V. & Schertl, H. P. (2008). Melting
of hydrous and carbonate mineral inclusions in garnet host during ultrahigh
pressure experiments. LITHOS 103, 25–45. EGU General Assembly, Vienna,
AUSTRIA, APR 24-29, 2005.
Plank, T. (2005). Constraints from Thorium/Lanthanum on sediment recycling
at subduction zones and the evolution of the continents. Journal of Petrology
46, 921–944.
Plank, T., Cooper, L. B. & Manning, C. E. (2009). Emerging geothermometers
for estimating slab surface temperatures. Nature Geoscience 2, 611–615.
Plank, T. & Langmuir, C. (1998). The chemical composition of subducting sed-
iment and its consequences for the crust and mantle. Chemical Geology 145,
325–394.
Poitrasson, F., Oelkers, E., Schott, J. & Montel, J. (2004). Experimental deter-
mination of synthetic NdPO4 monazite end-member solubility in water from
21 oC to 300 oC: Implications for rare earth element mobility in crustal fluids.
Geochimica et Cosmochimica Acta 68, 2207–2221.
Potts, P. J. (1996). A handbook of silicate rock analysis. London, UK: Chapman
and Hall.
Powell, R., Holland, T. & Worley, B. (1998). Calculating phase diagrams involv-
ing solid solutions via non-linear equations, with examples using THERMO-
CALC. Journal of Metamorphic Geology 16, 577–588.
Prowatke, S. & Klemme, S. (2005). Effect of melt composition on the parti-
tioning of trace elements between titanite and silicate melt. Geochimica et
Cosmochimica Acta 69, 695–709.
Pyle, J. & Spear, F. (2000). An empirical garnet (YAG) - xenotime thermometer.
Contributions to Mineralogy and Petrology 138, 51–58.
Pyle, J., Spear, F., Rudnick, R. & McDonough, W. (2001). Monazite-xenotime-
garnet equilibrium in metapelites and a new monazite-garnet thermometer.
Journal of Petrology 42, 2083–2107.
238
Ragozin, A. L., Liou, J. G., Shatsky, V. S. & Sobolev, N. V. (2009). The timing
of the retrograde partial melting in the Kumdy-Kol region (Kokchetav Massif,
Northern Kazakhstan). Lithos 109, 274–284.
Rapp, R., Ryerson, F. & Miller, C. (1987). Experimental-evidence bearing on
the stability of monazite during crustal anatexis. Geophysical Research Letters
14, 307–310.
Rapp, R. & Watson, E. (1986). Monazite solubility and dissolution kinetics -
implications for the thorium and light rare-earth chemistry of felsic magmas.
Contributions to Mineralogy and Petrology 94, 304–316.
Ravna, E. & Terry, M. (2004). Geothermobarometry of UHP and HP eclogites
and schists - an evaluation of equilibria among garnet-clinopyroxene-kyanite-
phengite-coesite/quartz. Journal of Metamorphic Geology 22, 579–592.
Reinecke, T. (1998). Prograde high- to ultrahigh-pressure metamorphism and ex-
humation of oceanic sediments at Lago di Cignana, Zermatt-Saas Zone, western
Alps. Lithos 42, 147–189.
Ren, M., Parker, D. & White, J. (2003). Partitioning of Sr, Ba, Rb, Y, and LREE
between plagioclase and peraluminous silicic magma. American Mineralogist
88, 1091–1103.
Roedder, E. (1984). Fluid inclusions, volume Reviews in Mineralogy 12. Miner-
alogical Society of America.
Rollinson, H. & Windley, B. (1980). Selective elemental depletion during meta-
morphism of Archean granulites, Scourie, NW Scotland. Contributions to Min-
eralogy and Petrology 72, 257–263.
Root, D., Hacker, B., Gans, P., Ducea, M., Eide, E. & Mosenfelder, J. (2005).
Discrete ultrahigh-pressure domains in the Western Gneiss Region, Norway:
implications for formation and exhumation. Journal of Metamorphic Geology
23, 45–61.
Rozen, O. (1971). Rifean in the Kokchetav massif. Proceedings of USSR Academy
of science 7, 102–104.
Rubatto, D. & Hermann, J. (2003). Zircon formation during fluid circulation
in eclogites (Monviso, Western Alps): Implications for Zr and Hf budget in
subduction zones. Geochimica Et Cosmochimica Acta 67, 2173 – 2187.
239
Rubatto, D. & Hermann, J. (2007a). Experimental zircon/melt and zircon/garnet
trace element partitioning and implications for the geochronology of crustal
rocks. Chemical Geology 241, 38–61.
Rubatto, D. & Hermann, J. (2007b). Zircon behaviour in deeply subducted rocks.
Elements 3, 31–35.
Rubatto, D., Hermann, J. & Buick, I. S. (2006). Temperature and bulk composi-
tion control on the growth of monazite and zircon during low-pressure anatexis
(Mount Stafford, central Australia). Journal Of Petrology 47, 1973 – 1996.
Rubatto, D., Regis, D., Hermann, J., Boston, K., Engi, M., Beltrando, M. &
McAlpine, S. R. B. (2011). Yo-yo subduction recorded by accessory minerals
in the Italian Western Alps. Nature Geoscience 4, 338–342.
Rubatto, D., Williams, I. S. & Buick, I. S. (2001). Zircon and monazite re-
sponse to prograde metamorphism in the Reynolds Range, central Australia.
Contributions To Mineralogy And Petrology 140, 458 – 468.
Rudnick, R., Barth, M., Horn, I. & McDonough, W. (2000). Rutile-bearing refrac-
tory eclogites: Missing link between continents and depleted mantle. Science
287, 278–281.
Rudnick, R., McLennan, S. & Taylor, S. (1985). Large ion lithophile elements in
rocks from high-pressure granulite facies terrains. Geochimica et Cosmochimica
Acta 49, 1645–1655.
Ryan, J., Morris, J., Tera, F., Leeman, W. & Tsvetkov, A. (1995). Cross-arc
geochemical variations in the Kurile arc as a function of slab depth. Science
270, 625–627.
Ryerson, F. & Watson, E. (1987). Rutile saturation in magmas - implications for
Ti-Nb-Ta depletion in island-arc basalts. Earth and Planetary Science Letters
86, 225–239.
Salters, V. & Stracke, A. (2004). Composition of the depleted mantle. Geochem-
istry Geophysics Geosystems 5.
Sawyer, E. (2008). Atlas of Migmatites, volume 9 of The Canadian Mineralogist
Special Publication. National Research Council (Canada) Research Press. Pp.
371.
240
Scherrer, N., Gnos, E. & Chopin, C. (2001). A retrograde monazite-forming re-
action in bearthite-bearing high-pressure rocks. Schweizerische Mineralogische
Und Petrographische Mitteilungen 81, 369–378.
Schmidt, M., Dardon, A., Chazot, G. & Vannucci, R. (2004). The dependence of
Nb and Ta rutile-melt partitioning on melt composition and Nb/Ta fraction-
ation during subduction processes. Earth and Planetary Science Letters 226,
415–432.
Schnetger, B. (1994). Partial melting during the evolution of the amphibolite-
facies to granulite-facies gneisses of the Ivrea Zone, Northern Italy. Chemical
Geology 113, 71–101.
Seydoux-Guillaume, A., Wirth, R., Heinrich, W. & Montel, J. (2002). Experi-
mental determination of Thorium partitioning between monazite and xenotime
using analytical electron microscopy and X-ray diffraction Rietveld analysis.
European Journal of Mineralogy 14, 869–878.
Shannon, R. (1976). Revised effective ionic-radii and systematic studies of inter-
atomic distances in halides and chalcogenides. Acta Crystallographica Section
A 32, 751–767.
Shatagin, K., Degtyarev, K. & Astakhantsev, O. (1999). Isotope composition of
Sr and Nd in granitoids of the Kokchetav massiv. Doklady Russ. Acad. Sci.
369, 525–528.
Shatagin, K., Degtyarev, K., Golubev, V., Astakhantsev, O. & Kuznetsov, N.
(2001). Vertical and lateral heterogeneity of the crust beneath Northern Kaza-
khstan from geochronological and isotopic-geochemical data on Paleozoic gran-
itoids. Geotectonics 5, 26–44. (in Russian).
Shatsky, V., Jagoutz, E., Sobolev, N., Kozmenko, O., Parkhomenko, V. &
Troesch, M. (1999). Geochemistry and age of ultrahigh pressure metamor-
phic rocks from the Kokchetav massif (Northern Kazakhstan). Contributions
to Mineralogy and Petrology 137, 185–205.
Shatsky, V., Ragozin, A. & Sobolev, N. (2006a). Some aspects of metamorphic
evolution of ultrahigh-pressure calc-silicate rocks of the Kokchetav Massif. Rus-
sian Geology and Geophysics 47, 105–119.
241
Shatsky, V., Sitnikova, E., Koz’menko, O., Palessky, S., Nikolaeva, I. & Zay-
achkovsky, A. (2006b). Behavior of incompatible elements during ultrahigh-
pressure metamorphism (by the example of rocks of the Kokchetav massif).
Russian Geology and Geophysics 47, 482–496.
Shatsky, V., Sobolev, N. & Vavilov, M. (1995). Diamond bearing metamorphic
rocks of the Kokchetav Massif (northern Kazakhstan). In: R. Coleman &
X. Wang (eds.) Ultrahigh pressure metamorphism, pages 427–455. Cambridge:
Cambridge University Press.
Shatsky, V. S., Sobolev, N. V., Korsakov, A. V., Ragozin, A. L. & Zayachkovsky,
A. A. (2005). A new occurrence of diamondiferous rocks in Kokchetav massif
(Northern Kazakhstan). In: Proceedings of Eclogite conference 2005.
Sitnikova, E. S. & Shatsky, V. S. (2009). New FTIR spectroscopy data on the
composition of the medium of diamond crystallization in metamorphic rocks
of the Kokchetav Massif. Russian Geology and Geophysics 50, 842–849.
Skora, S. & Blundy, J. (2010). High-pressure hydrous phase relations of radio-
larian clay and implications for the involvement of subducted sediment in arc
magmatism. Journal of Petrology 51, 2211–2243.
Smith, D. (1980). A tectonic melange of foreign eclogites and ultramafites in west
Norway. Nature 287, 366–367.
Sobolev, N., Shatskii, V., Vavilov, M. & Goryainov, S. (1991). Coesite inclusionin zircon from diamond-containing gneisses of Kokchetav massif - 1st find ofcoesite in metamorphic rocks in the USSR. Doklady Akademii Nauk SSSR
321, 184–188.
Sobolev, N., Shatskii, V., Vavilov, M. & Goryainov, S. (1994). Zircon fromhigh-pressure metamorphic rocks of folded regions as an unique container ofinclusions of diamond, coesite and coexisting minerals. Doklady Akademii Nauk
334, 488–492.
Sobolev, N. & Shatsky, V. (1990). Diamond inclusions in garnets from metamor-
phic rocks - a new environment for diamond formation. Nature 343, 742–746.
Sobolev, N. V., Schertl, H. P., Valley, J. W., Page, F. Z., Kita, N. T., Spicuzza,
M. J., Neuser, R. D. & Logvinova, A. M. (2011). Oxygen isotope variations
of garnets and clinopyroxenes in a layered diamondiferous calcsilicate rock
242
from Kokchetav Massif, Kazakhstan: a window into the geochemical nature
of deeply subducted UHPM rocks. Contributions to Mineralogy and Petrology
162, 1079–1092.
Spandler, C., Hermann, J., Arculus, R. & Mavrogenes, J. (2003). Redistribution
of trace elements during prograde metamorphism from lawsonite blueschist to
eclogite facies; implications for deep subduction-zone processes. Contributions
To Mineralogy And Petrology 146, 205 – 222.
Spandler, C., Mavrogenes, J. & Hermann, J. (2007). Experimental constraints on
element mobility from subducted sediments using high-P synthetic fluid/melt
inclusions. Chemical Geology 239, 228 – 249.
Spear, F. S. & Pyle, J. M. (2010). Theoretical modeling of monazite growth in a
low-Ca metapelite. Chemical Geology 273, 111–119.
Stockhert, B., Duyster, J., Trepmann, C. & Massonne, H. (2001). Microdiamond
daughter crystals precipitated from supercritical COH plus silicate fluids in-
cluded in garnet, Erzgebirge, Germany. Geology 29, 391–394.
Stoeckhert, B., Trepmann, C. A. & Massonne, H. J. (2009). Decrepitated UHP
fluid inclusions: about diverse phase assemblages and extreme decompression
rates (Erzgebirge, Germany). Journal of Metamorphic Geology 27, 673–684.
Syracuse, E. M., van Keken, P. E. & Abers, G. A. (2010). The global range of
subduction zone thermal models. Physics of the Earth and Planetary Interiors
183, 73–90.
Tailby, N. D., Walker, A. M., Berry, A. J., Hermann, J., Evans, K. A., Mavro-
genes, J. A., O’Neill, H. S. C., Rodina, I. S., Soldatov, A. V., Rubatto, D. &
Sutton, S. R. (2011). Ti site occupancy in zircon. Geochimica et Cosmochimica
Acta 75, 905–921.
Taylor, M. & Ewing, R. C. (1978). The crystal structures of the ThSiO4 poly-
morphs: huttonite and thorite. Acta Crystallographica Section B 34, 1074–
1079.
Taylor, S. & McLennan, S. (1985). The Continental Crust: its composition and
evolution. Palo Alto, California: Blackwell Scientific Pub.
Theunissen, K., Dobretsov, N. L., Korsakov, A., Travin, A., Shatsky, V. S.,
Smirnova, L. & Boven, A. (2000). Two contrasting petrotectonic domains in
243
the Kokchetav megamelange (north Kazakhstan): Difference in exhumation
mechanisms of ultrahigh-pressure crustal rocks, or a result of subsequent de-
formation? Island Arc 9, 284 – 303.
Tomkins, H. S., Powell, R. & Ellis, D. J. (2007). The pressure dependence of the
zirconium-in-rutile thermometer. Journal of Metamorphic Geology 25, 703–
713.
Tropper, P., Manning, C. E. & Harlov, D. E. (2011). Solubility of CePO4 mon-
azite and Y PO4 xenotime in H2O and H2O−NaCl at 800 oC and 1 GPa: Im-
plications for REE and Y transport during high-grade metamorphism. Chem-
ical Geology 282, 58 – 66.
Turkina, O. M., Letnikov, F. A. & Levin, A. V. (2011). Mesoproterozoic Grani-
toids of the Kokchetav Microcontinent Basement. Doklady Earth Sciences 436,
176–180.
van Keken, P., Kiefer, B. & Peacock, S. (2002). High-resolution models of subduc-
tion zones: Implications for mineral dehydration reactions and the transport
of water into the deep mantle. Geochemistry Geophysics Geosystems 3.
Van Orman, J., Grove, T., Shimizu, N. & Layne, G. (2002). Rare earth ele-
ment diffusion in a natural pyrope single crystal at 2.8 GPa. Contributions to
Mineralogy and Petrology 142, 416–424.
Veksler, I., Dorfman, A., Kamenetsky, M., Dulski, P. & Dingwell, D. (2005). Par-
titioning of lanthanides and Y between immiscible silicate and fluoride melts,
fluorite and cryolite and the origin of the lanthanide tetrad effect in igneous
rocks. Geochimica et Cosmochimica Acta 69, 2847–2860.
Volkova, N. I., Tarasova, E. N., Polyanskii, N. V., Vladimirov, A. G. &
Khomyakov, V. D. (2008). High-pressure rocks in the serpentinite melange
of the Chara zone, Eastern Kazakhstan: Geochemistry, petrology, and age.
Geochemistry International 46, 386–401.
Wang, C. Y., Campbell, I. H., Stepanov, A. S., Allen, C. M. & Burtsev, I. N.
(2011). Growth rate of the preserved continental crust: II. Constraints from
Hf and O isotopes in detrital zircons from Greater Russian Rivers. Geochimica
et Cosmochimica Acta 75, 1308–1345.
244
Wang, Q. & Cong, B. (1999). Exhumation of UHP terranes: A case study from
the Dabie Mountains, eastern China. International Geology Review 41, 994–
1004.
Ward, C., McArthur, J. & Walsh, J. (1992). Rare-earth element behavior during
evolution and alteration of the Dartmoor Granite, SW England. Journal of
Petrology 33, 785–815.
Wark, D. & Miller, C. (1993). Accessory mineral behavior during differentiation
of a granite suite - monazite, xenotime and zircon in the Sweetwater Wash
Pluton, Southeastern California, USA. Chemical Geology 110, 49–67.
Watson, E. & Harrison, T. (1983). Zircon saturation revisited: temperature and
composition effects in a variety of crustal magma types. Earth and Planetary
Science Letters 64, 295–304.
Watson, E. & Harrison, T. (1984). Accessory minerals and the geochemical evolu-
tion of crustal magmatic systems - a summary and prospectus of experimental
approaches. Physics of the Earth and Planetary Interiors 35, 19–30.
Watson, E. B. & Mueller, T. (2009). Non-equilibrium isotopic and elemental frac-
tionation during diffusion-controlled crystal growth under static and dynamic
conditions. Chemical Geology 267, 111–124.
Watson, E. B., Wark, D. A. & Thomas, J. B. (2006). Crystallization thermome-
ters for zircon and rutile. Contributions to Mineralogy and Petrology 151, 413
– 433.
White, J., Holt, G., Parker, D. & Ren, M. (2003). Trace-element partitioning
between alkali feldspar and peralkalic quartz trachyte to rhyolite magma. Part
I: Systematics of trace-element partitioning. American Mineralogist 88, 316–
329.
White, T. & Dong, Z. (2003). Structural derivation and crystal chemistry of
apatites. Acta Crystallographica Section B-Structural Science 59, 1–16.
Williams, I. (1998). U-Th-Pb geochronology by ion microprobe. In: W. M. McK-
ibben (ed.) Applications of Microanalytical Techniques to Understanding Min-
eralizing Process, volume 7 of The Economic Geology, Reviews in Economic
Geology, chapter 1, pages 1–36. Society of Economic Geologists, Inc.
245
XiaoYing, G., ShuNing, L. & YongFei, Z. (2011). On the study multiphase solid
inclusions in UHP metamorphic minerals. Acta Petrologica Sinica 27, 469–489.
Xiong, X., Keppler, H., Audetat, A., Ni, H., Sun, W. & Li, Y. (2011). Partitioning
of Nb and Ta between rutile and felsic melt and the fractionation of Nb/Ta
during partial melting of hydrous metabasalt. Geochimica et Cosmochimica
Acta 75, 1673–1692.
Yang, P. & Pattison, D. (2006). Genesis of monazite and Y zoning in garnet from
the Black Hills, South Dakota. Lithos 88, 233–253.
Yasnygina, T. A. & Rasskazov, S. V. (2008). Tetrad effect in rare earth element
distribution patterns: Evidence from the Paleozoic granitoids of the Oka zone,
Eastern Sayan. Geochemistry International 46, 814–825.
Yurimoto, H., Duke, E., Papike, J. & Shearer, C. (1990). Are discontinuous
chondrite-normalized ree patterns in pegmatitic granite systems the results of
monazite fractionation. Geochimica et Cosmochimica Acta 54, 2141–2145.
Zhang, R., Liou, J., Ernst, W., Coleman, R., Sobolev, N. & Shatsky, V. (1997).
Metamorphic evolution of diamond-bearing and associated rocks from the
Kokchetav Massif, northern Kazakhstan. Journal of Metamorphic Geology
15, 479–496.
Zhang, R. Y., Liou, J. G. & Ernst, W. G. (2009). The Dabie-Sulu continental
collision zone: A comprehensive review. Gondwana Research 16, 1–26.
Zhao, Z.-F., Zheng, Y.-F., Chen, R.-X., Xia, Q.-X. & Wu, Y.-B. (2007). Ele-
ment mobility in mafic and felsic ultrahigh-pressure metamorphic rocks during
continental collision. Geochimica et Cosmochimica Acta 71, 5244–5266.
Zheng, Y.-F., Gao, X.-Y., Chen, R.-X. & Gao, T. (2011a). Zr-in-rutile ther-
mometry of eclogite in the dabie orogen: constraints on rutile growth during
continental subduction-zone metamorphism. Journal of Asian Earth Sciences
40, 427–451.
Zheng, Y.-F., Xia, Q.-X., Chen, R.-X. & Gao, X.-Y. (2011b). Partial melting,
fluid supercriticality and element mobility in ultrahigh-pressure metamorphic
rocks during continental collision. Earth-Science Reviews 107, 342 – 374.
246
Zhimulev, F. I. (2007). The Tectonics and Early Ordovician Geodynamic Evo-
lution of the Kokchetav HP-UHP Metamorphic Belt. Ph.D. thesis, IGM SO
RAN, Novosibirsk.
Zhimulev, F. I., Buslov, M. M., Travin, A. V., Dmitrieva, N. V. & De Grave,
J. (2011). Early-Middle Ordovician nappe tectonics of the junction between
the Kokchetav HP-UHP metamorphic belt and the Stepnyak paleoisland arc
(northern Kazakhstan). Russian Geology and Geophysics 52, 109–123.
Zhimulev, F. I., Poltaranina, M. A., Korsakov, A. V., Buslov, M. M., Druzyaka,
N. V. & Travin, A. V. (2010). Eclogites of the Late Cambrian-Early Ordovician
North Kokchetav tectonic zone (northern Kazakhstan): structural position and
petrology. Russian Geology and Geophysics 51, 190–203.
Zhou, H. (2007). Quantitative Geochemistry. Imperial College Press, 304 pages.
Zhou, J.-B., Wilde, S. A., Zhao, G.-C., Zheng, C.-Q., Jin, W., Zhang, X.-Z. &
Cheng, H. (2008). Detrital zircon U-Pb dating of low-grade metamorphic rocks
in the Sulu UHP belt: evidence for overthrusting of the North China Craton
onto the South China Craton during continental subduction. Journal of the
Geological Society 165, 423–433.
Zhu, X. & O’Nions, R. (1999). Zonation of monazite in metamorphic rocks and
its implications for high temperature thermochronology: a case study from the
Lewisian terrain. Earth and Planetary Science Letters 171, 209–220.
Zindler, A. & Hart, S. (1986). Chemical geodynamics. Annual Review of Earth
and Planetary Sciences 14, 493–571.
Zonenshain, L., Kuzmin, M. & Natapov, L. (1990). Geology of the USSR: a
plate-tectonic synthesis . Geodynamics Serie. AGU, 242 pages.