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Zircon
Tiny but Timely
LEFT: Zircon grain (400 µm)
crystallized from a 530 Ma
high-temperature partial
melt, Prydz Bay,
Antarctica. PHOTOMICROGRAPH (CROSSED POLARIZERS)
BY NIGEL KELLY
Simon L. Harley and Nigel M. Kelly*
W
here would Earth science be without zircon? Tiny crystals of zircon
occur in many rocks, and because their atomic structure remains
stable over very long periods of geological time, they are able to
provide a picture of the early history of the Earth and of the evolution of the
crust and mantle. Zircon has long been recognized as the best geochronometer
using the radioactive decay of uranium to lead. Recent developments in analytical
techniques, using small-diameter laser, ion and electron beams, high-precision
mass spectrometry and a variety of microscopic imaging methods, allow us to
obtain the ages of tiny volumes of complex crystals that record stages in their
long growth history. Coupled measurements of the isotopes of oxygen and
hafnium provide a mineralogical window into the separation of the Earth’s
crust from the mantle and the temperature and character of processes involved
in crustal evolution. Zircon is being used to unravel ever more complex
geological systems, presenting exciting opportunities for research on this
remarkable mineral.
in the crust, zircon potentially contains a record, in its oxygen isotope
composition, of the role of lowtemperature versus high-temperature processes in defining the character of source regions for melts
(Valley 2003). Measurements of
the U, Th and He contents of zircon
can be used to infer the rates at
which recently active landscapes
developed and the times at which
the exposed rocks cooled to nearsurface temperatures. Finally, zircon
offers significant potential as a
phase into which the highly
radioactive isotopic products of
nuclear reactions may be introduced in order to evaluate their
impact on mineral structures (e.g.
Hanchar and van Westrenen 2007).
KEYWORDS: zircon, geochronology, continental crust, U–Pb, Hf, trace elements
Central to all of these applications
is the behaviour of zircon in complex Earth systems. Thanks to important developments in secondary ion mass spectrometry
(SIMS), laser- ablation induced coupled plasma – mass spectrometry (LA-ICP–MS) and low-blank thermal-ionisation mass
spectrometry (TIMS), we are now able to accurately and precisely measure a battery of useful trace element and isotopic
signatures in zircon. The interpretation of these isotopic
and compositional data in terms of ages, isotopic reservoirs
and processes requires the careful and systematic integration of microanalysis with petrology and mineral characterisation.
INTRODUCTION
Zircon, ZrSiO4, is a mineral of singular importance in Earth
science. Its widespread use in geochronology, based on the
decay of uranium (U) to lead (Pb), has established it as
Earth’s timekeeper. Thus, zircon records the ages of hallmark events in Earth history, including its earliest evolution,
the oldest sediments, extinction episodes, mountain-building
events, and supercontinent construction and dispersal
(Rubatto and Hermann 2007 this issue; Harley et al. 2007
this issue).
Recent developments in microanalysis have extended the
range of Earth problems that can be addressed using zircon.
As a phase enriched in hafnium (Hf) compared to radioactive lutetium (Lu), zircon retains a strong fingerprint of the
isotopic character of the sources of the magmatic rocks in
which it crystallizes, evidence that is critical for models of
formation and growth of continental crust (Scherer et al.
2007 this issue). As a phase that can accommodate significant
amounts of temperature- or process-sensitive trace elements,
including the rare earth elements (REE, or lanthanides),
yttrium (Y) and titanium (Ti), zircon can also provide compelling chemical evidence for the mineral–melt–fluid
processes operating during crust formation and maturation,
hydrothermal alteration and diagenesis (Hanchar and van
Westrenen 2007 this issue; Harley et al. 2007; Geisler et al.
2007 this issue). Despite the ravages of cycling through and
This issue of Elements focuses on the advances in our understanding of zircon and highlights the gaps in our knowledge that have emerged from in situ isotopic, chemical,
spectroscopic and microtextural studies on zircon formed at
high and low temperatures. The systems range from meltbearing ultrahigh-temperature metamorphic environments
(Harley et al. 2007) and ultrahigh-pressure metamorphic
environments (Rubatto and Hermann 2007) to lowertemperature hydrothermal environments in which zircon
behaviour may be dictated by its response to aggressive fluids (Geisler et al. 2007). Understanding the processes that
operate in each environment is critical to the interpretation
of zircon age data and the hafnium isotope information
used to infer continental growth rates and earliest Earth history (Scherer et al. 2007). The importance of this understanding is highlighted by new experiments and models
relating to the partitioning of trace elements among zircon,
melts and fluids and to the dissolution–reprecipitation of
zircon in fluids (Hanchar and van Westrenen 2007; Geisler
et al. 2007).
* Grant Institute of Earth Science, The University of Edinburgh
Edinburgh EH9 3JW, United Kingdom
E-mail: [email protected]
[email protected]
ELEMENTS, VOL. 3,
PP.
13–18
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F EBRUARY 2007
There is still much to learn about zircon and its behaviour.
Zircon cannot be treated simply as a passive ‘safehouse’ of
stored isotopic and chemical information but must instead
be interpreted carefully, in its petrological, mineralogical
and geological contexts, and in the light of all possible lines
of evidence. Zircon has been a wonderful servant in our
quest to unravel the history of the Earth but has much more
to offer as we unlock the secrets of its chemical and physical
responses to Earth processes.
WHAT IS ZIRCON ANYWAY?
Zircon is a tetragonal orthosilicate mineral in which isolated SiO4 tetrahedra are linked through sharing their edges
and corners with intervening ZrO8 dodecahedra (FIG. 1).
These ZrO8 dodecahedra share edges to form zigzag chains
along the b axis, whereas along the c axis, edges are shared
with the SiO4 tetrahedra to produce chains with alternating
SiO4 and ZrO8 polyhedra. These sets of chains are separated
by channels or voids that are unoccupied in pure zircon.
However, in natural zircon, these channels may contain
interstitial impurities at parts per million (ppm) to tens of
ppm levels (Hoskin et al. 2000; Hanchar et al. 2001). The
c-axis chains are important in controlling the anisotropic
physical properties of zircon and its common prismatic
habit. The effect of Zr4+−Si4+ repulsion on the symmetry of
the cation sites in zircon is considerable. First, the SiO4
tetrahedra are elongated by 13% along the c axis compared
with their size along the a axis. Second, the ZrO8 polyhedra
resolve into two interpenetrating ZrO4 tetrahedra (FIG. 1).
One type of tetrahedron has longer Zr−O bond lengths
(0.227 nm) parallel to the c axis and can be visualised as
forming elongate chains by sharing two of its edges with
the alternating SiO4 tetrahedra (FIG. 2). The other type of
ZrO4 tetrahedron has shorter Zr−O bond lengths (0.213 nm)
and shares its four corners with SiO4 tetrahedra. Further
details of the crystal structure of zircon are presented in the
excellent review by Finch and Hanchar (2003).
A view of the zircon structure projected from the a axis,
stripped of the cross-linking shortened ZrO4 tetrahedra to
highlight the chains of edge-sharing, alternating SiO4–ZrO4 tetrahedra
parallel to the c axis (after Finch and Hanchar 2003). SiO4 tetrahedra are
coloured yellow, and elongated ZrO4 tetrahedra blue.
FIGURE 2
The overall structure of zircon is relatively open, with unoccupied space represented by the channels parallel to the
c axis and the void volumes bounded by SiO4 and ZrO8
polyhedra, as shown in FIGURE 1. This structure results in
zircon’s moderately high density, 4.66 g cm-3, and also contributes significantly to its very low absolute thermal
expansion and compressibility and to the anisotropy of
these parameters. Most of the small amount of volume
expansion in zircon, about 0.6% from room temperature to
its stability limit of 1690°C, is accommodated along the c
axis by increase in the Zr−O bond length. Although pure
zircon is highly incompressible, it is anisotropic in its compressibility because the Zr−O bonds parallel to the c axis are
able to shorten preferentially. These properties make pure
zircon or zircon with low contents of trace elements
extremely resistant to physical modification related to
changes in pressure or temperature, and render it an excellent refractory mineral that is potentially useful for storage
of radioactive and toxic isotopes provided they have suitable ionic radii.
WHAT IS ZIRCON GOOD FOR?
Even with its underlying structural beauty, zircon would be
of little interest to most Earth scientists if it were just plain
ZrSiO4. However, the simplicity of zircon’s chemical formula belies its great scope for chemical diversity at the trace
to minor element level. An indication of the variety of uses
of zircon is provided in TABLE 1. Since most cations in zircon
have very low diffusivity, many of zircon’s chemical signatures are preserved either from the time of its formation or
from the last significant geological process to have acted on
and modified its chemistry (Cherniak and Watson 2003).
A view of the zircon structure projected from the a axis
onto the plane defined by the b and c axes (after Finch
and Hanchar 2003). SiO4 tetrahedra are coloured yellow, and ZrO8
dodecahedra are in shades of blue. One dodecahedron is unshaded to
reveal its sub-structure consisting of two distorted ZrO4 tetrahedra. The
elongated ZrO4 tetrahedra share upper and lower edges with the SiO4
tetrahedra.
FIGURE 1
ELEMENTS
Zircon can incorporate many elements, e.g. P, Sc, Nb, Hf, Ti,
U, Th and REE, in trace (up to thousands of ppm) or minor
(up to 3 wt%) amounts. These elements are incorporated
through a number of single-site and coupled-cation substitution mechanisms (Hoskin and Schaltegger 2003). The pri14
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TABLE 1
KEY CHEMICAL FEATURES OF ZIRCON AND THEIR APPLICATIONS
Chemical/physical
property
Substitutions/
other points to note
Key applications
Comments
U and Th
U up to 5000 ppm
Th up to 1000 ppm
(U4+, Th4+) = Si4+
U–Pb geochronology
A concordia diagram can be used to evaluate the isotopes
of U and their Pb decay products. Th/U ratios, used in the
past to distinguish magmatic from metamorphic and
hydrothermal zircon, must be treated with caution.
He
Formed by decay
of U and Th
Determination of exhumation and
landscape development rates using
U–Th–He thermochronometry
Low-temperature chronometry is based on the closure of
zircon to He loss. This method gives an age related to the
time the zircon cooled through ~40°C.
Hf
HfO2 mostly <3 wt%
Hf4+ = Si4+
Ti
Ti up to 120 ppm
Ti4+ = Si4+
Y and REE
Y mostly <5000 ppm
Total REE <2500 ppm
(Y3+, REE3+)P5+ = Zr4+Si4+
O isotope composition
176Lu decays to 176Hf. High Hf/Lu in zircon means its ratio
Investigation of crustal residency and
continental growth; crustal versus mantle
of 176Hf/177Hf changes very little with time, so it can be
sources of magmas in which zircon formed used to infer sources by reference to an Earth model.
Ti is maximised when zircon is in equilibrium with rutile.
Ti thermometry can yield T of zircon crystallization, which
Ti in zircon thermometry
usually occurs late in the cooling of a magma, or T of
metamorphic zircon growth with rutile.
Reconstruction of magmatic histories;
fingerprinting of magma sources;
tuning of ages to mineral reactions
Requires extensive knowledge of trace element partitioning
among zircon, melts and competitor minerals over a range
of P, T, composition and oxygen fugacity conditions.
Fingerprinting the contribution of
sediments and crust to the sources of
magmas; examining crustal recycling
Significant fractionation of 18O from 16O occurs at low T.
Variations in 18O/16O isotopic composition of zircon are
used to determine the role of sources that have been
affected by low-T fractionation.
zircon formation and by competition between zircon and
other minerals in which some of these elements may be
more compatible (Hanchar and van Westrenen 2007;
Harley et al. 2007). REE and Y contents will be affected also
by the operation of other coupled substitutions and by the
requirement to charge-balance additional cations, such as
Mg, Fe, Ca and Al, that may be incorporated on the interstitial sites described above (Hoskin et al. 2000; Geisler et al.
2007). Compilations of zircon from high- and low-temperature geological environments and formed through a variety of processes show that Y contents generally range
between 10 and 5000 ppm and total REE typically between
100 and 2500 ppm (e.g. Hoskin and Schaltegger 2003).
Considerable discussion has arisen as to the use of REE patterns and abundances as discriminants of the provenance of
detrital zircon populations (Hoskin and Ireland 2000) and,
in the case of zircon xenocrysts, of source-rock lithologies
(Belousova et al. 2002). Considerations surrounding these
and other uses of zircon REE data are explored further by
Hanchar and van Westrenen (2007).
mary controls on the substitutions are the ionic radii of the
substituting cations compared with Zr4+ and Si4+ cations.
Substitutions that minimise strain effects on either or both
sites will be favoured. The crystal-chemical limitations are
that Zr4+ in 8-fold co-ordination has an ionic radius of
0.084 nm and Si4+ in tetrahedral co-ordination an ionic
radius of 0.026 nm. On this basis it has been suggested that
(OH)4 can replace SiO4. There can be considerable substitution of Hf4+ (ionic radius = 0.083 nm) on the 8-fold Zr4+ site,
and a solid solution towards the mineral hafnon (HfSiO4)
exists. Zircon generally contains considerable HfO2
(TABLE 1), which is central to its utility as an indicator of
crustal residence and growth via Hf isotope analysis
(Hawkesworth and Kemp 2006; Scherer et al. 2007). U4+
(8-fold ionic radius 0.10 nm), Th4+ (0.105 nm) and Ti4+
(0.074 nm) can also be accommodated, generally at much
lower abundance levels, on this site. Although U can reach
wt% levels, its concentration is usually less than 5000 ppm,
while the abundances of Th (<1000 ppm) and Ti (<120 ppm)
are lower still. The incorporation of Ti into zircon, particularly if it has formed in equilibrium with rutile and quartz,
is temperature sensitive and provides the basis for a new zircon geothermometer (Watson et al. 2006). This may be
used, with caution, to determine the temperature of crystallization of magmas, migmatites and zircon–rutile assemblages in metamorphic rocks. Because of its 8-fold ionic
radius of 0.129 nm, Pb2+ is not incorporated into growing
zircon crystals at more than ppb levels under most conditions, which is crucial to geochronology.
The cation substitutions mentioned above often lead to the
production of spectacular internal textures in zircon that
may be interpreted in terms of growth histories and diffusion−reaction or dissolution–reprecipitation processes.
These features, for example oscillatory and sector zoning,
can be imaged using cathodoluminescence (CL), backscattered electron imagery and a variety of other techniques
(e.g. forescatter imagery, Nomarski interference imagery,
infrared spectrometry, atomic force microscopy, Raman
spectroscopy). Yet another of the many remarkable properties of zircon is that it often contains negligible amounts of
cations that suppress its cathodoluminescence response,
and so it can be imaged to very high resolution using CL. In
addition, most of the substitutions involve replacement of
the lower mass elements Zr and Si by much heavier elements (U, Th, REE, Hf), resulting in major shifts in average
atomic mass, which are well imaged by changes in backscattered electron intensity.
The most important coupled substitution involving both
the Zr4+ and Si4+ sites in zircon is that commonly referred
to as the ‘xenotime’ substitution. This substitution (TABLE 1)
involves Y and REE substituting for Zr and charge-balancing
P5+ (4-fold ionic radius = 0.029 nm) substituting for Si,
which would ultimately produce xenotime, (Y,REE)PO4.
Scandium (Sc3+, with an ionic radius of 0.087 nm) also substitutes for Zr, in quantities up to 250 ppm, in a similar coupled substitution. In principle the heavy trivalent REE, with
their smaller ionic radii (e.g. 8-fold Lu3+ = 0.0977 nm), Y3+
(8-fold ionic radius 0.1019 nm) and tetravalent Ce4+ (8-fold
ionic radius 0.097 nm) are closer to Zr4+ and so will be more
favourably incorporated in the zircon structure (i.e. they are
more compatible) than the larger light trivalent REE (e.g.
La3+ = 0.116 nm). However, the absolute and relative abundances of the REE and Y in zircon will be influenced by
their abundances and availability in the environments of
ELEMENTS
At this point it is worth recalling that zircon is an orthosilicate mineral, with the Zr and Si cations and their substitutes bound to oxygen anions in tetrahedral and 8-fold
co-ordination. Given the slow diffusivities inferred and documented for oxygen in zircon (Cherniak and Watson 2003;
Valley 2003), it is likely that magmatic zircon can preserve
an oxygen isotope composition that was in equilibrium
with the magma from which it crystallized. In confirmation
15
F EBRUARY 2007
of this, magmatic zircon grains with inherited xenocrystic
cores have in several cases been shown to have distinctive
18O/16O isotopic compositions (expressed as δ18O in per mil
units, ‰, relative to the Vienna Standard Mean Ocean
Water – VSMOW), interpreted as reflecting differences in
the isotopic compositions of the magmas in which they
formed. Given that mantle-derived zircon has a δ18O value
of +5.5‰, and that O isotopes are significantly fractionated
at low temperatures to produce potential crustal source
rocks (e.g. sediments) with much higher oxygen isotope
ratios (i.e. δ18O values of +12 to +14‰), the δ18O values of
granitoid zircon may be used to trace the involvement of
older, chemically evolved crust in magma genesis (Valley
2003). In other words, the oxygen isotope ratio of zircon
can be used to discriminate between new, mantle-derived
crust and crust that has been reworked (Valley 2003). This
approach is even more powerful when combined with zircon U–Pb age data and Hf isotope information on the same
analysed grains (e.g. Hawkesworth and Kemp 2006) These
considerations highlight the importance of zircon as a
geochronometer, the principles of which will be explained
in the remaining section of this introduction.
of the zircon grains from any one of the three ‘clocks’, by
measuring the appropriate isotopic ratio and solving for
time using the relevant exponential equation. We can go
further and obtain three age estimates if we measure all
three isotopic ratios. In an ideal closed system, the three age
estimates would agree within the errors of measurement.
However, in real zircon grains, we cannot assume that Th
and U are equally ‘closed’ to post-crystallization effects. It is
also essential to correct for any Pb initially present prior to
the accumulation of radiogenic Pb in the grain, as this
inherited Pb would lead to erroneously old age estimates.
Yet another of the wonderful features of crystallizing zircon
is that it can only incorporate negligible amounts of ambient
Pb into its structure. Nearly all of the Pb in zircon is produced by U and Th decay – so corrections for inherited Pb
are in most instances very small relative to the amounts of
true radiogenic Pb* present in the grains. Inherited (or
‘common’) Pb can be corrected for by analysing nonradiogenic 204Pb, where present in the zircon, and then subtracting the 206Pb and 207Pb that would be associated with
this 204Pb at a chosen reference age – the ‘common Pb
correction’.
The usual approach to zircon geochronology is to consider
the U–Pb system alone, as there is no natural non-nuclear
means of fractionating 235U from 238U. In addition, as the
modern-day ratio of 235U/238U is well known (1/137.88),
the need to actually analyse very low abundances of 235U is
obviated [i.e. 207Pb*/235U = 137.88(206Pb*/238U)]. Instead of
solving the 235U and 238U decay systems separately, we can
plot mutually compatible sets of daughter/parent ratios,
207Pb*/235U and 206Pb*/238U, that would evolve in the zircon
grains as time elapses since their formation (i.e. as they age).
This is the basis of the concordia diagram (FIG. 3; Wetherill
1956). As this article is being written, we are celebrating the
50th anniversary of this elegant graphical device, which has
opened the door to the systematic assessment of zircon and
other accessory mineral U–Pb isotopic data.
BACK TO BASICS:
ZIRCON U–PB GEOCHRONOLOGY
The potential of zircon as a mineral geochronometer was
recognised by Holmes (1911), amongst others, well before
the isotopes of Pb could be measured and long before 235U
was identified as the second radioactive isotope of uranium.
It is now known that there are three distinct radioactive
decay series involving the parent isotopes 238U, 235U and
232Th, which produce as their final daughter products the
isotopes 206Pb, 207Pb and 208Pb, respectively. Each of these
decay processes involves several intermediate steps and
short-lived intermediate isotopes. For example, the decay of
238U to 206Pb occurs via a chain of intervening alpha-decay
steps (liberating 4He α-particles) coupled with beta-decay
steps (releasing a β-particle and transforming a neutron to a
proton). These steps yield short-lived isotopes that decay in
seconds, years, decades or hundreds of thousands of years.
However, because the final step in the decay series is many
orders of magnitude slower than the earlier steps, the whole
decay process can be mathematically described by a single
decay equation relating the number of ultimate parent
atoms remaining (e.g. 238U) and the number of final radiogenic daughter atoms (e.g. 206Pb*) to time:
206Pb*/238U
238t
= eλ
–1
The concordia curve itself is the locus of the mutually compatible or concordant 207Pb*/235U and 206Pb*/238U ratios,
both of which increase outwards from the origin as the time
since zircon crystallization passes by. At time zero, when
(1)
where e is the exponential function, t is time, and λ is the
decay constant specific to this decay scheme, i.e. λ238 =
1.55125e-10. 206Pb* refers to the radiogenic 206Pb accumulated in the crystal as a result of the decay of 238U. In the
case of 238U, it takes approximately 4468 million years for
half the 238U initially present in the grain to decay to 206Pb
– this is the half-life, which again is characteristic of the specific decay scheme. Similar expressions can be formulated
for 207Pb* produced from the decay of 235U and 208Pb* produced from 232Th, with λ235 = 9.8485e-10 and λ232 = 4.9475e-11.
The half-life of 235U is 704 million years and that of 232Th
about 14 billion years.
In the following we will consider the case of a suite of zircon
grains formed at a single time, perhaps through crystallization of a felsic magma. Through the incorporation of U and
Th at the time of growth, in every zircon grain there will, in
effect, be three different ‘clocks’ ticking away, each at its
characteristic rate. By virtue of the half-lives noted above,
235U decays about 7 times faster than 238U does, while 232Th
decays even more slowly. In principle, we could find the age
ELEMENTS
Concordia diagram. Note that the axes are the ratios
of the radiogenic daughter Pb isotopes divided by their
respective parent U isotope. The concordia curve traces out the compatible ratios as they develop with time elapsed since the moment the
‘clocks’ are set in newly formed zircon, and is annotated with this time
elapsed, or ‘age’. Lines a and b show the evolution of the compatible
ratios in 1000 and 3000 Ma zircon respectively (see text).
FIGURE 3
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the zircon forms, there is no radiogenic Pb* in the zircon.
Some 1000 million years later, these ratios have increased to
1.6777 and 0.1678, respectively (FIG. 3, line a), because the
two types of radiogenic Pb have accumulated through
decay of their respective parent U isotope. So, if we sampled
a rock formed 1000 Ma ago and analysed its zircon grains,
we would, ideally, expect them to preserve these
207Pb*/235U and 206Pb*/238U ratios and hence lie on the
concordia curve at the point corresponding to 1000 Ma. By
the same logic, if we analysed zircon grains from an undisturbed rock that is now 3000 Ma old, we would expect the
207Pb*/235U and 206Pb*/238U ratios to have evolved to
18.1988 and 0.5926, respectively (FIG. 3 line b). These are
concordant zircon grains: the two Pb*/U ratios as measured
in the grains correspond to the same age (time elapsed), and
the position along concordia is a direct measure of age. The
individual analysis points plot as elongated ellipses or polyhedra because the two ratios, and hence their errors, are
strongly correlated.
A
The elegance of the concordia diagram is that it also allows
you to investigate and evaluate zircon Pb*/U analyses that
do not lie on concordia. In such cases, where the ages determined from the 207Pb*/235U and 206Pb*/238U ratios do not
agree, the zircon grains are said to be discordant. They are
normally discordant if they lie below the concordia curve,
which is the usual type of discordancy, and are termed
reverse discordant if the analyses lie above the concordia
curve (i.e. at a higher 206Pb*/238U for a given 207Pb*/235U ratio).
B
An example of the former situation is illustrated on the concordia diagrams of FIGURE 4. In this case discordant zircon
Pb*/U analyses have been generated through the superposition
of a second, real, geological event on a suite of zircon grains
some 3000 million years after they initially formed. This
event not only formed new zircon crystals, possibly with
distinctive morphologies and trace element signatures, but
also caused variable loss (without any Pb isotope fractionation) of the previously accumulated radiogenic Pb from
domains in the older zircon grains (FIG. 4A). Whatever the
precise mechanism causing this Pb loss (e.g. Geisler et al.
2007), the large ionic radius of Pb2+ renders it susceptible to
removal from the old zircon grains.
Concordia diagram and episodic Pb loss. Variable loss of
Pb from older zircon grains, arising because of the effects
of a second event, produces a discordia between the older formation
age (now 4000 Ma) and the age of the second event (1000 Ma).
FIGURE 4
FIGURE 4B illustrates the effect on the concordia diagram if
the zircon grains plotted in FIG. 4A evolve for a further 1000
million years. All grains will now have accumulated radiogenic Pb* over this 1000 million years, so that those newly
formed in the second event will be concordant at 1000 Ma
(FIG. 4B). If all the radiogenic Pb* accumulated up to the
time of the second event was lost from the first-generation
zircon grains, then these would be completely reset, and all
memory of the antiquity of these 4000 Ma old grains would
be lost; they too would record ages of 1000 Ma. Alternatively,
and more typically, the loss of Pb in these old zircon grains
is only partial or confined to spaced microdomains. In this
case a series of discordant analyses may result, spread out
along a discordia (FIG. 4B) that intersects the concordia at
two points: an upper (older) intercept giving the formation
age of the zircon grains – 4000 Ma in this case – and a lower
intercept giving an age of 1000 Ma, in accord with the
newly formed grains.
of cores and overgrowths in such cases may yield concordant populations that define the ages of the growth/modification events, but even with careful microtextural control,
it is not uncommon to find that some zircon populations
define discordia.
Complex U–Pb isotopic data can be interpreted and utilised
to gain insights into Earth history, but only in concert with
complementary chemical and isotopic information, which
includes the REE, Ti, Hf isotopes and O isotopes, briefly
alluded to above and listed in TABLE 1. A number of these
key signatures of zircon behaviour and the processes involved
in zircon modification and recrystallization are considered
in the articles that follow.
The case described above is an idealised example of episodic
Pb loss, which is a major cause of discordant zircon populations. Nevertheless, it serves to illustrate how the concordia
diagram can be ‘read’ to reveal several pieces of geologically
useful information. As will be shown in the articles that follow, we now are very aware through microanalytical and
mictotextural studies that many zircon grains (especially in
metamorphic rocks, but also in granites) preserve more
than one period of growth or grain modification. Analysis
ELEMENTS
ACKNOWLEDGMENTS
We thank the editors of Elements for providing us with this
forum for presenting some of the current developments in
zircon research. In particular, we thank Ian Parsons and
Pierrette Tremblay for their tireless editorial and organisational work on our behalf. We also thank all contributors
and reviewers, without which this issue would not have
been possible. .
17
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GLOSSARY
Accessory mineral – A mineral that represents <1% of the modal pro-
LA-ICP–MS (laser-ablation inductively coupled plasma – mass
portion of the minerals in a rock. These minerals are commonly (but
not in all cases) enriched in trace elements important for the interpretation of rock history.
spectrometry) – A microanalytical method that employs a focused
laser beam to ablate material from samples. The ejected matter is
ionised using a plasma before being passed through to a mass spectrometer.
Backscattered electron imaging – A technique in which the highenergy electrons produced by elastic collision of an incident electron
beam with the electron cone of a sample atom are imaged using a
detector within an SEM. The yield of backscattered electrons is proportional to the number of sample electrons, therefore offering the
potential to image chemical zoning within a target sample.
Monazite – A phosphate mineral rich in light rare earth elements (LREE)
BSE (bulk silicate Earth) – What is left after subtracting the core from
and lithosphere through collision of continents and/or accretion of
microcontinents or island arcs to a continent at the end of a Wilson
Cycle.
and commonly Th, based on the general formula (La,Ce,Th)PO4. It
incorporates appreciable U and Y (up to several weight %) through
coupled substitutions.
Orogenesis – Mountain building resulting from thickening of the crust
the bulk Earth composition; also referred to as ‘primitive mantle’.
Cathodoluminescence (CL) – Light emitted from minerals during bombardment by an electron beam, typically imaged using a detector fitted to a scanning electron microscope or an optical microscope. Luminescence within minerals is enhanced by particular elements (e.g. Dy)
but is potentially suppressed by other elements (e.g. U, Th, Fe) or
defects in the crystal structure. Zoning observed in CL images therefore reflects chemical variations within individual crystals.
Supracrustal rocks – Sedimentary or volcanic rocks formed at or near
CHUR (chondritic uniform reservoir) – CHUR’s characteristics are
P–T paths – When rock masses are buried deep in the Earth’s crust, the
derived from meteorite studies. The BSE and CHUR reservoirs are
assumed to have identical ratios of refractory lithophile elements (e.g.
Lu/Hf, Sm/Nd).
‘paths’ along which they move can be described in terms of the
changes in pressure (depth) and temperature. The relationship
between pressure and temperature over the time span of the metamorphism reflects the tectonic history of the rocks in question.
Concordance – Agreement of ages obtained by the
235U–207Pb
238U–206Pb
Earth’s surface. Those found today in ancient portions of continents
were buried and metamorphosed before being exposed by uplift and
erosion. They provide important clues (e.g. basalt pillows, cross bedding) about ancient depositional environments and contain the only
records of early life.
and
chronometers.
Recrystallization – A general term used to refer to in situ modification
or change in the character, and potentially the chemistry, of a mineral
or rock. Recrystallization may be partial or total, and may involve distinct processes such as reaction–diffusion or replacement, and coupled
dissolution–reprecipitation.
Depleted Mantle – Mantle that has had melt extracted from it. Some
elements fractionate more strongly into the melt phase. As a result, the
depleted mantle’s Lu/Hf and Sm/Nd values are higher than those of BSE.
238U–206Pb and 235U–207Pb ages.
This commonly results from the partial loss of radiogenic Pb by a zircon during alteration or heating and recrystallization. For recent Pb
loss, the ratio of the Pb daughter isotopes may be used to calculate a
207Pb/206Pb age, which may still date zircon crystallization. If the Pb
loss is ancient, the 207Pb/206Pb age is only a minimum crystallization age.
Discordance – Disagreement between
SIMS (secondary ion mass spectrometry) – Also referred to as an ion
microprobe or microscope, SIMS measures the chemical or isotopic
composition of small sample volumes by focusing a beam of highenergy primary ions onto a polished sample surface, ablating atoms
and molecules, and generating secondary ions that are analysed by
mass spectrometry. The high spatial resolution offered by SIMS (commonly <30 µm wide and <1 µm deep during analysis of geological
materials) allows in situ analysis of geological materials. The method is
relatively non-destructive, allowing multiple analyses to be performed
within single grains or zones within grains, but has lower analytical
precision than ID-TIMS.
ID-TIMS (isotope-dilution thermal-ionisation mass spectrometry)
– A method of isotopic analysis in which an isotopic tracer is added to
a dissolved sample (e.g. zircon) to make a homogeneous isotopic mixture,
the isotopic composition of which is analysed using TIMS. This technique is currently the form of isotopic measurement with the highest
precision, but it requires complete dissolution of part of or whole grains.
Zoning – Variation in chemical composition among different domains
Inheritance – The presence of older zircon grains in an igneous or meta-
within a single crystal. The geometry of zoning patterns preserved in
crystals may give insights into the processes that led to the formation
of new crystals or the alteration of previously existing grains. For example, oscillatory zoning (regular, finely spaced, circular zoning patterns)
is commonly preserved in magmatic zircon crystals and reflects fluctuating trace element concentrations in the parent magma.
morphic rock that did not crystallize from that rock’s parent magma.
The inherited zircon grains may have been incorporated into a magma
through the partial melting of a pre-existing zircon-bearing rock or through
assimilation of zircon-bearing country rocks during magma ascent.
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F EBRUARY 2007