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Transcript
U SERIES DISEQUILIBRIA: INSIGHTS INTO MANTLE
MELTING AND THE TIMESCALES OF MAGMA
DIFFERENTIATION
David W. Peate
Department of Geoscience
University of Iowa
Iowa City, Iowa, USA
Chris J. Hawkesworth
Department of Earth Sciences
University of Bristol
Bristol, UK
Received 22 March 2004; revised 18 December 2004; accepted 17 January 2005; published 31 March 2005.
[1] Several U series nuclides have half-lives (230Th, 76 kyr;
231
Pa, 33 kyr; and 226Ra, 1.6 kyr) comparable to timescales
of magmatic processes. We review the basic principles of
extracting time information from U series nuclides and
summarize variations in (230Th/238U), (226Ra/230Th), and
(231Pa/235U) observed in magmas from mid-ocean ridges,
within-plate settings, and subduction zones to contrast
melt generation processes in different tectonic settings. U
series disequilibria on melt and crystal phases of igneous
rocks can provide temporal information on different
stages in the magmatic history (melting duration, melt
transport rates, magmatic crustal residence times, and
timing of crystal growth) and potentially provide clues
about the nature and mineralogy of mantle sources,
mantle upwelling rates and porosity, fluid influences, and
mechanisms of melt generation and transport. The subject
is beginning to take a genuinely integrated approach to
developing physically realistic quantitative models that
offer increasingly exciting opportunities in the study of
magmatic processes.
Citation: Peate, D. W., and C. J. Hawkesworth (2005), U series disequilibria: Insights into mantle melting and the timescales of magma
differentiation, Rev. Geophys., 43, RG1003, doi:10.1029/2004RG000154.
1.
INTRODUCTION
[2] In order to develop quantitative models for magmatic processes it is necessary to have constraints on the
likely rates and timescales of processes such as melting,
differentiation, and crystallization, and yet such information has been difficult to estimate from the geological
record. However, the increasingly widespread application
of U series isotopes, driven by advances in analytical
methods in the late 1980s, has revolutionized our understanding of melt generation and crystallization processes
and their timescales. Unlike many radioactive nuclides that
decay directly to a stable daughter nuclide (e.g., 87Rb !
87
Sr, 147Sm ! 143Nd, and 176Lu ! 176Hf ), the long-lived
isotopes of uranium and thorium (238U, 235U, and 232Th)
decay ultimately to isotopes of Pb (206Pb, 207Pb, and
208
Pb, respectively) via a series of intermediate unstable
nuclides that have half-lives ranging from 250 kyr (thousand years) to microseconds (Figure 1). Attention has
focused mainly on the decay products of 238U and 235U,
the so-called U series nuclides, because several of these
intermediate nuclides have half-lives (t1/2) (230Th, t1/2
76 kyr; 231Pa, t1/2 33 kyr; 226Ra, and t1/2 1.6 kyr; see
Table 1) that are comparable with the inferred timescales
of magmatic processes.
[3] The utility of these U series nuclides stems from
the unique properties of the decay chains. If the system is
undisturbed, then it is said to be in ‘‘secular equilibrium’’
in which the activity of all nuclides in the decay chain
are equal. The activity of a nuclide is the rate of
radioactive decay of that nuclide in the sample. Typical
measurement units for activity are decays per minute per
gram (dpm g1) or Bequerels per gram (Bq g1 or
decays per second per gram). The activity of a nuclide
is simply the product of the number of atoms of the
nuclide and the decay constant l, where l = ln(2)/t1/2. As
the parent nuclide at the top of each decay chain has a
much longer half-life than all the intermediate nuclides,
the activity or flux of radioactive decay from this parent
nuclide remains essentially constant over timescales relevant to the study of the shorter-lived intermediate
nuclides. Because different elements have distinct chemical properties, it is possible for magmatic processes to
change the relative abundances of the different U series
nuclides from the proportions expected for secular equilibrium, such that the activities of the nuclides in the
decay chain are no longer equal. This elemental fractionation leads to disequilibrium between parent and daughter
nuclides within the chain, and the return to secular
equilibrium is governed by the half-life of the shorterlived nuclide. It is this process of reestablishing secular
equilibrium that allows timescales of magmatic processes
to be determined. Disequilibrium between a parent and
daughter nuclide pair is detectable up to about five half-
Copyright 2005 by the American Geophysical Union.
Reviews of Geophysics, 43, RG1003 / 2005
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Paper number 2004RG000154
8755-1209/05/2004RG000154$15.00
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Peate and Hawkesworth: URANIUM SERIES DISEQUILIBRIA
Figure 1. Periodic table summary of parts of the 238U, 235U, and 232Th decay chains. Nuclides of
interest are outlined in black with the half-lives given below. The very short-lived intermediate nuclides
of Pa and Th are indicated as shaded. The dotted arrows indicate the decay of 226Ra, 231Pa, and 228Ra to
206
Pb, 207Pb, and 208Pb, respectively, via a series of very short-lived intermediate nuclides that are
omitted for clarity.
lives of the shorter-lived nuclide, depending on the initial
extent of disequilibrium and the precision of the measurements. For example, if 230Th-238U disequilibrium is measured in a sample (i.e., activity of 230Th does not equal
activity of 238U), then this indicates that an event that
chemically fractionated U from Th took place less than
380 kyr ago as the half-life of 230Th is 76 kyr.
[4] Processes such as partial melting, crystal growth,
and fractional crystallization can all potentially produce
disequilibrium between U series nuclides with different
chemical properties. Therefore we should, in principle, be
able to use these nuclides to constrain transport rates of
melts from mantle sources, the residence time of magma in
crustal reservoirs, and the timing of crystal growth in
addition to estimating eruption ages. During partial melting, disequilibria between U series nuclides will depend
critically on the source mineralogy, the rate of melting, and
the degree of interaction with the surrounding mantle as
the melt percolates upward. Therefore U series disequilibria data on primitive magmas can potentially also provide
important information about the nature and mineralogy of
the mantle source, the porosity during melting, and
plausible mechanisms of melt generation and transport
in the mantle.
[5] The 230Th-231Pa-226Ra isotopes are the most widely
used in the study of igneous rocks, but there is also
considerable potential to exploit isotopes with even shorter
half-lives, including 228Ra (t1/2 5.75 years), 210Pb (t1/2
22 years), and 210Po (t1/2 138 days), particularly to look at
timescales of magma degassing and the timing of very recent
crystal growth [see Condomines et al., 2003, and references
therein]. This paper aims to provide an accessible and up-todate review for a broad audience of the applications of U
series disequilibria to magmatic processes, focusing on the
230
Th-238U, 226Ra-230Th, and 231Pa-235U schemes. In the
decade or so since the last published general reviews [e.g.,
Gill and Condomines, 1992; Gill et al., 1992; Macdougall,
1995], significant advances have been made, particularly
through detailed comprehensive regional studies combining
230
Th-226Ra-231Pa data on otherwise well-characterized
samples and more advanced modeling incorporating new
ideas and observations from other areas of geophysics and
geochemistry.
[6] This paper is structured in four parts: (1) a summary
of the basic principles, concepts, and assumptions behind
U series disequilibria applications; (2) a summary of the
degree and sense of disequilibria observed between
230
Th-238U, 226Ra-230Th, and 231Pa-235U in young igneous
rocks from different tectonic environments; (3) a discussion of how disequilibria are used to constrain timescales
of magma differentiation and crystal growth; and (4) a
discussion of the controlling factors in generating the
observed disequilibria in magmas from different tectonic
settings, together with implications for melt generation
models. More detailed recent reviews on these individual
TABLE 1. Half-Lives and Decay Constants for U Series
Nuclides of Interesta
Nuclide
232
Th
U
235
U
234
U
230
Th
231
Pa
226
Ra
238
a
2 of 43
Half-Life, years
Decay Constant (l), year1
14.010 ± 0.014 109
4.4683 ± 0.0048 109
0.7038 ± 0.0010 109
245,250 ± 490
75,690 ± 230
32,760 ± 220
1,599 ± 4
4.9475 ± 0.0050 1011
1.5513 ± 0.0017 1010
9.849 ± 0.013 1010
2.8263 ± 0.0057 106
9.158 ± 0.028 106
2.116 ± 0.014 105
4.335 ± 0.011 104
Recommended values are from Bourdon et al. [2003a].
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Peate and Hawkesworth: URANIUM SERIES DISEQUILIBRIA
Figure 2. Disequilibrium within each of the parentdaughter nuclide pairs (226Ra-230Th, 231Pa-235U, and
230
Th-238U) returning to secular equilibrium (i.e., activity
ratio equal to 1) over a timescale governed by the half-life
of the daughter nuclide (i.e., 226Ra, 231Pa, and 230Th,
respectively). Figure 2 also emphasizes that U series
disequilibrium methods are only applicable to geologically
very young samples. Examples are given for the temporal
evolution of disequilibria in samples with initial daughter
excesses of 20% and initial daughter deficits of 20%.
topics are available [see Bourdon et al., 2003c; Rudnick,
2003].
2. BACKGROUND TO U SERIES DISEQUILIBRIA
IN IGNEOUS ROCKS
2.1. Systematics and Conventions
[7] Isotope ratios of the U series nuclides are usually
expressed as activity ratios, and by convention they are
enclosed in parentheses, e.g., (230Th/238U). Activity ratios
are easily related to the atomic ratio by the ratio of the
respective decay constants, i.e., ( 230 Th/ 238 U) activity =
[230Th/238U]atomic [l230/l238]. At secular equilibrium
the activities of all nuclides in a particular decay chain are
equal, and so ( 230 Th/ 238 U) = 1, ( 226 Ra/ 230 Th) = 1,
(231Pa/235U) = 1, and (234U/238U) = 1. As fresh igneous
rocks are both expected and usually observed to have
(234U/238U) = 1, then 238U can essentially be treated as
the parent of 230Th. It is important to remember that for all
these U series nuclide pairs both the parent and daughter
nuclides are unstable and will undergo radioactive decay at
different rates depending on their respective half-lives.
[8] If a sample has measured disequilibrium between a
particular parent-daughter pair because of some chemical
process that affected the two elements to different extents, it
is said to have an excess of the nuclide with the higher
activity. For example, a mid-ocean ridge basalt with
(230Th/238U) = 1.20 has a 20% 230Th excess. In this case,
there is an insufficient quantity of the parent nuclide 238U to
maintain this amount of excess daughter 230Th through
radioactive decay, and the excess 230Th is therefore said
to be ‘‘unsupported.’’ This excess will decay away at a rate
governed by the 230Th half-life until the amount of 230Th is
equal to that being produced by radioactive decay from the
238
U in the sample.
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[9] On the other hand, a subduction zone lava with
(230Th/238U) = 0.80 has a 20% deficit of the daughter
nuclide, and the amount of 230Th is less than that being
produced from decay of 238U. In this case the amount of
230
Th will gradually increase at a rate determined by the
230
Th half-life until the amount of 230Th in the sample is
equal to that being produced by 238U decay.
[10] As it takes about five half-lives of the shorter-lived
nuclide for secular equilibrium to be restored between a
parent and daughter nuclide pair, this means that the
different U series nuclide pairs provide information on
different timescales (Figure 2): The 230Th-238U system
records U-Th fractionation within the last 380 kyr, the
226
Ra-230Th system records Ra-Th fractionation within the
last 8 kyr, and the 231Pa-235U system records Pa-U fractionation within the last 165 kyr. Note that igneous geochemists use the word ‘‘fractionation’’ in a broad sense to
cover the effects of any process that changes the relative
abundances of different elements or isotopes. The magnitude of 230Th-238U, 226Ra-230Th, and 231Pa-235U disequilibria measured in a sample today will be governed by several
factors: (1) the extent of chemical fractionation between
parent and daughter nuclides (e.g., partial melting, source
mineralogy, crystallization, diffusion, melt immiscibility,
and melt-matrix interaction); (2) the elapse of time since
the fractionation event during which radioactive decay will
occur (e.g., melt transport time from mantle source regions
to the crust, residence time in crustal magma reservoirs, and
time since lava eruption); and (3) surface weathering or
hydrothermal alteration processes.
[11] It should now be apparent that U series disequilibria
methods can essentially be applied only to geologically very
young samples (Figure 2): For example, samples older that
about 380 ka should have (238U/230Th) = 1, as secular
equilibrium will have been reestablished. Furthermore, the
eruption ages of samples need to be well known in order to
be able to correct for any posteruptive radioactive decay. For
example, if it is known only that a sample was erupted
sometime in the last 8000 years that is fine for the 238U-230Th
system because this is a short time period compared to the
76 kyr 230Th half-life. However, for the 226Ra-230Th
system, which will return to secular equilibrium over a
similar period of about 8000 years, it is clear that the
eruption ages would need to be known more precisely.
2.2. Analytical Developments
[12] The most significant analytical advance (see reviews
by Chen et al. [1992] and Goldstein and Stirling [2003]) has
been the development of thermal ionization mass spectrometric (TIMS) techniques for the measurement of 234U
[Chen et al., 1986], 230Th [Edwards et al., 1987; Goldstein
et al., 1989; Palacz et al., 1992], 226Ra [Cohen and
O’Nions, 1991; Volpe et al., 1991; Chabaux et al., 1994],
and 231Pa [Pickett et al., 1994; Bourdon et al., 1999a].
Earlier studies used a spectrometry, which counts the
number of a particles emitted by radioactive decay from a
sample in a given time. For nuclides such as 230Th that have
relatively slow decay rates and low abundances in rocks,
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long counting times on the order of tens of weeks are
needed to achieve acceptable counting statistics. Mass
spectrometry, on the other hand, counts the number of ions
rather than waiting for the radioactive decay of the nuclides,
thus achieving far more favorable counting statistics for the
nuclides of interest here (234U, 235U, 238U, 230Th, 232Th,
226
Ra, and 231Pa). The 2s precisions of 0.1– 2% can be
obtained by mass spectrometer measurements of less than
2 hours duration, compared with 2 – 10% for much longer
a spectrometry measurements, and sample size requirements for mass spectrometry are also significantly less than
for a spectrometry (e.g., U-Th, 0.1 – 2 mg versus 1 – 100 mg,
and Pa-Ra, 0.003-1 pg versus 10– 50 pg) [Goldstein and
Stirling, 2003]. This reduction in sample size has been
especially advantageous for studies looking at magma
chamber processes and geochronology by analysis of separated mineral phases from rocks.
[13] Mass spectrometric measurement of the isotopic
composition of thorium poses several particular challenges.
First, most igneous rocks have Th/U ratios of 3 ± 1, which is
equivalent to an atomic 230Th/232Th ratio of (6 ± 3) 106.
Thus for silicate samples, there is a huge contrast (5 orders
of magnitude) in beam intensity between the major 232Th
peak and the minor 230Th peak. Collisions with residual
molecules within the mass spectrometer cause ion scattering, leading to a low mass tail from the large 232Th peak that
partly obscures the small 230Th peak. An energy filter is
needed to reduce this background from scattered ions so as
to improve the detectability of the 230Th peak. Second,
thorium is a difficult element to ionize thermally, with less
than 1% of the sample ionized. Therefore the recent
development of multicollector inductively coupled plasma
mass spectrometers, which are much more efficient at
ionizing Th and other elements, offer much potential
for the future, particularly in further reducing sample
size requirements and improving measurement precision.
Plasma-based techniques have been developed for the
measurement of 230Th/232Th and 234U/238U [Luo et al.,
1997; Turner et al., 2001b; Shen et al., 2002], 226Ra
[Pietruszka et al., 2002], and 231Pa [Regelous et al., 2004].
2.3. Robust Indicators of Magmatic Processes?
[14] If U series disequilibria are to be used to provide
information about magmatic processes, it is important to
confirm that the observed disequilibria are primary and
magmatic in origin and that they do not result from
posteruptive alteration through interactions with seawater
or groundwater or from shallow level contamination with
sedimentary or hydrothermal materials. Most concern has
centered on mid-ocean ridge basalts (MORB), where the
potential for contamination is high given the relatively low
initial elemental concentrations and the availability of
materials potentially enriched in various U series nuclides.
Palagonite rinds and crystalline pillow lava interiors have
elevated uranium contents relative to unaltered glass
samples, indicating pervasive uptake of uranium from
seawater [e.g., Macdougall et al., 1979]. Fe-Mn oxides
and hydrothermal precipitates that often coat MORB
RG1003
glasses contain significant amounts of unsupported 230Th,
231
Pa, and 234U scavenged from seawater, and inclusion of
even small amounts of such material will significantly
influence the measured disequilibria. Therefore careful
handpicking of glass chips free of any visible alteration
and Fe-Mn oxide coatings, together with chemical leaching
to remove material in cracks, is critical for analysis of
MORB glasses. Bourdon et al. [2000a] used the abundance
of 10Be to place limits on the contribution of metalliferous
sediments to the measured U series disequilibria in handpicked and leached MORB glasses. Beryllium 10 is a
radioactive nuclide (t1/2 1.5 Myr) that is only produced
at significant rates through cosmogenic processes in the
atmosphere. It is highly enriched in the metalliferous
sediments found near the ridge axes, but it can be assumed
to be absent in the pristine mantle-derived glasses, and
thus any 10Be found in the MORB glass must result from
near-surface sedimentary contamination. The 10Be data
demonstrate that sedimentary contamination can only
account for <1% of the measured excess 230Th and
231
Pa in carefully selected and leached samples.
[15] Although 226Ra is not strongly concentrated in
Fe-Mn oxides, it is likely to be highly enriched in hydrothermal Ba-rich phases such as barite, which is a ubiquitous
mineral in the ridge environment [Volpe and Goldstein,
1993]. It is more difficult to assess how the measured
226
Ra-230Th disequilibria in individual samples might have
been influenced by assimilation of even minor amounts of
barite, although one would expect elevated Ba contents
relative to other highly incompatible elements such as Rb
or Th. The near constancy of Ba/Rb for both mid-ocean ridge
basalts and ocean island basalts and the limited variations of
Ba/Th with 226Ra excesses suggest that barite assimilation
has a minimal influence on the composition of ridge
basalts [Lundstrom et al., 1999; Elliott and Spiegelman,
2003].
[16] The 234U/238U ratio of a sample can provide a useful
monitor of alteration by seawater or groundwaters. At
magmatic temperatures, 234U is not expected to be fractionated from 238U, and so fresh igneous rocks should have
(234U/238U) = 1, whereas most natural waters show disequilibrium between 234U and 238U, principally because of alpha
recoil effects. The physical process of alpha decay by a 238U
atom results in recoil of the daughter nuclide 234Th (which
rapidly decays to 234U) because of conservation of momentum, so that it is displaced from its original location. It is
either ejected directly into the surrounding fluid phase or
sits in a now damaged site in the crystal where it is more
susceptible to mobilization by fluids. Seawater is well
mixed with respect to uranium and has (234U/238U) of
1.146 [Chen et al., 1986].
[17] The effects of seawater alteration are illustrated in
Figure 3 using two different sample suites: (1) submarine
glass and crystalline pillow interior powders from a single
MORB sample from the Reykjanes Ridge [Peate et al.,
2001a; D. Peate, unpublished data, 1996] and (2) subaerial
arc lavas from a single coastal outcrop on Miyakejima
volcano, Japan [Yokoyama et al., 2003]. In each case the
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Peate and Hawkesworth: URANIUM SERIES DISEQUILIBRIA
Figure 3. (234U/238U) versus (232Th/238U) diagram showing the effects of seawater contamination on coastal
subaerial arc lavas (Miyakejima volcano, Japan [Yokoyama
et al., 2003]) and submarine mid-ocean ridge basalts
(Reykjanes Ridge [Peate et al., 2001a; D. Peate,
unpublished data, 1996). Altered samples (shaded symbols) lie on a linear mixing line between fresh samples
with (234U/238U) = 1 (solid symbols) and seawater with
(234U/238U) = 1.146 and (232Th/238U) = 0 (solid diamond).
Note that even minor seawater contamination that elevates
(234U/238U) in a sample by just 1% will produce a
corresponding decrease of almost 7% in the Th/U ratio
relative to the pristine sample, and it will have an effect
of similar magnitude on the 230Th-238U disequilibrium
[Yokoyama et al., 2003].
variably altered samples lie on a linear mixing line between
pristine samples with (234U/238U) = 1 and seawater, and an
elevation in (234U/238U) of just 1% corresponds to almost
7% difference in (232Th/238U) from the ‘‘true’’ pristine
value. This observation highlights why it is critical to
measure (234U/238U) to a high precision 5% [e.g., Sims
et al., 2002a; Yokoyama et al., 2003] on all subaerial
samples from coastal regions and on submarine samples
in order to screen for altered samples. However, (234U/238U)
data are often not published along with the other disequilibria data (226Ra-230Th-238U and 231Pa-235U), and it is
notable that less than 40% of the data compiled for the
global review of disequilibria in igneous rocks discussed in
section 3 have accompanying (234U/238U) data.
2.4. Principal Mechanisms to Produce 238U-230Th,
Ra-230Th, and 231Pa-235U Disequilibria
2.4.1. Trace Element Partitioning
[18] We can define trace elements as those elements that
have low abundances in igneous rocks, i.e., less than about
0.1% or 1000 ppm (parts per million) by weight. This
definition includes all of the U series nuclides of interest
to us (i.e., U, Th, Ra, and Pa). The concept of the partition
coefficient (or distribution coefficient) is an important tool
to describe quantitatively the behavior of trace elements
in magmatic systems. The partition coefficient simply
226
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measures the relative partitioning of a trace element
between two phases. The Nernst partition coefficient (D)
is used to quantify the partitioning of a trace element at
equilibrium, usually between a mineral and a melt phase.
It is given by the concentration of the element in the
mineral divided by the concentration of the element in the
melt. For example, if a clinopyroxene phenocryst contains
1 ppm Th and the glassy matrix of the host lava contains
20 ppm Th, then DTh clinopyroxene/melt = 0.05. When more
than one mineral phase is present, such as during mantle
melting (olivine + orthopyroxene + clinopyroxene + garnet
or spinel), a bulk partition coefficient must be calculated to
describe partitioning of a trace element between the melt
and the bulk solid. This is done by weighting the mineral/
melt partition coefficients of the individual mineral phases
by their proportions in the bulk solid.
[19] If an element has a mineral/melt partition coefficient
of less than 1, this means that the element prefers to be in
the melt rather than in that particular mineral, and it is said
to be incompatible in the mineral. Conversely, a partition
coefficient greater than 1 implies that the element preferentially partitions into the mineral relative to the melt, and it is
said to be compatible in the mineral. Partition coefficients
are a measure of the extent to which a trace element can be
accommodated in the crystal structure of a particular mineral, which will be governed by the ionic charge and radius
of the incorporated element and the nature of the crystallographic sites within the mineral. The values of specific
mineral/melt partition coefficients have to be determined
experimentally, and they vary as a function of temperature,
pressure, and the composition of the mineral. However,
crystal lattice strain models are being developed to allow
extrapolation of the experimental results to other temperatures and pressures and compositions and even to predict
values for other elements (see Blundy and Wood [2003] for
a comprehensive review of partitioning data for U series
nuclides).
2.4.2. Crystallization
[20] As a magma cools, different mineral phases nucleate
and begin to grow. The different minerals will incorporate
or exclude different trace elements to different degrees,
depending on the value of the relevant mineral/melt partition coefficient. If one element is preferentially incorporated
into a certain mineral phase relative to another element, then
removal of this mineral phase from the magma during
fractional crystallization will change the ratio of the two
elements in the magma, i.e., lead to fractionation of the two
elements. The extent to which this elemental ratio is
changed will depend on the ratio of the partition coefficents
of the two elements, the absolute values of the partition
coefficients, and the amount of the mineral removed from
the magma.
[21] U and Th are highly incompatible elements in most
common crystallizing minerals (olivine, pyroxene, feldspar,
amphibole, mica, and Fe-Ti oxides) with partition coefficients <0.05 [Blundy and Wood, 2003]. Thus these common
phases will all have very low concentrations of U and Th
relative to the host magma. The low U and Th concentrations
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in the crystalling phases mean that not much U or Th will
be extracted from the magma as the crystals settle out
during magmatic differentiation. Therefore closed system
fractional crystallization does not usually result in significant variations in the U/Th ratios of related whole rock
samples, irrespective of any differences in the relative
partition coefficients for U and Th. Fractional crystallization is therefore also unlikely to be a significant cause of
238
U-230Th disequilibrium in basaltic magmas. However,
238
U-230Th disequilibrium can be produced in the crystallizing minerals because small differences in partition
coefficients will determine the relative extents to which
238
U and 230Th are incorporated into the crystals from the
melt, even if these amounts are small compared to the
abundances in the melt.
[22] Crystallization of small amounts of accessory phases
with very high U and/or Th contents (e.g., zircon, apatite,
sphene, chevkinite, allanite, and monazite), on the other
hand, can cause significant variations in U/Th ratios in more
differentiated magmas. For example, phonolitic glasses
from the 13 ka Laacher See zoned eruption in Germany
show a 12% variation in U/Th that Bourdon et al. [1994]
interpreted as being due to rapid crystallization and removal
of apatite and sphene, which are accessory phases with
high-Th and high-U contents and low U/Th.
[23] Radium partitions into feldspar crystals more readily
than Th (DTh 0.003), and it becomes a compatible element
(i.e., DRa > 1) for many alkali feldspar compositions [Blundy
and Wood, 2003]. Feldspar crystals (especially alkali feldspar) will develop elevated 226Ra/230Th values compared to
the magma, and thus their removal during differentiation
should lower 226Ra/230Th values in the magma. Fractional
crystallization of feldspar is therefore a potential cause of
226
Ra-230Th disequilibrium in more evolved magmas. The
effect of fractional crystallization on 231Pa-235U disequilibrium is uncertain because of a lack of experimentally
determined partition coefficient data and a lack of direct
measurements of 231Pa-235U in separated minerals. Both
Pa and U are likely to be highly incompatible in most
common crystallizing minerals [Blundy and Wood, 2003],
and so minimal effect on 231Pa-235U disequilibrium during
fractional crystallization is expected. However, Pa might
be partitioned strongly into oxides (e.g., ilmenite and
rutile) and zircon [Blundy and Wood, 2003], and so these
phases might influence 231Pa-235U disequilibrium during
differentiation of more evolved magmas.
2.4.3. Melt Generation and Transport
[24] The melting process is probably the principal
mechanism that produces 238U-230Th, 226Ra-230Th, and
231
Pa-235U disequilibrium in magmas. Disequilibrium during melting originates through differences in the chemical
behavior of the various U series nuclides, as reflected
primarily in different partition coefficients. A common
assumption is made that the source rocks are in secular
equilibrium, which will be the case unless there has been
any recent chemical modification of the source. Therefore
the initial activity ratios of the source will be known, i.e.,
(238U/230Th) = 1, (226Ra/230Th) = 1, and (231Pa/235U) = 1,
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which is not the case for trace element ratios. This
assumption might not hold in certain situations: for example, if melting in subduction zones is triggered by recent
addition of a fluid rich in U and Ra but not Th and Pa.
[25] The simplest melting models treat the U series
nuclides like any other trace elements and attribute any
disequilibrium between U series nuclides simply to net
elemental fractionation as controlled by differences in
partitioning between melt and residual solid. They include
a range of different physical models of melt generation from
‘‘batch melting’’ (where all the melt produced stays in
equilibrium with the residual solids until it is extracted
instantaneously to the surface) to ‘‘fractional melting’’
(where infinitesimal amounts of melt are extracted instantaneously as soon as they are produced), but the common
feature is that they are ‘‘time-independent.’’ A characteristic
of any equilibrium melting model is that significant trace
element fractionation will only occur when the degree of
melting is similar to or less than the bulk partition coefficients. For example, bulk partition coefficients for U and
Th in typical peridotitic mantle will be 103, and those for
Ra and Pa are likely to be at least a factor of 10 smaller
[Blundy and Wood, 2003]. In this case the degree of melting
would have to be extremely small, on the order of a percent
or less, in order to produce significant disequilibrium
between the U series nuclides.
[26] Melting takes place over a finite timescale that is
likely to be similar in magnitude to the half-lives of 230Th,
231
Pa, and 226Ra. Thus the duration of melt generation and
melt extraction will also play a critical role in determining
the degrees of disequilibrium between these nuclides that
are preserved in a magma exiting the melting region. This
will be influenced by the physical dynamics of the melt
generation process in terms of parameters such as melting
rate, porosity, melt velocity, solid upwelling rate, and the
length of the melting column. Melting models that explicitly
take into account the radioactive ingrowth and decay of the
daughter nuclides during melt production are often referred
to as ‘‘ingrowth’’ models.
[27] In these models, movement of melt relative to the
solid during the melting process (i.e., two-phase flow
[McKenzie, 1985]) will lead to disequilibrium between the
U series nuclides because of differences in the residence
time of parent and daughter nuclides within the melting
region. To explain the principle of daughter ingrowth in
simple terms, consider the slow melting of a garnet peridotite mantle source. For this composition, Th is more
incompatible than U (i.e., DTh < DU), and so the 230Th that
was initially in secular equilibrium with 238U in the source
will be preferentially partitioned into the first melt formed,
relative to 238U. This leaves the residual solid deficient in
230
Th (i.e., (230Th/238U) <1), and additional 230Th will begin
to accumulate (or ingrow) from decay of the 238U in the
solid as the system attempts to reestablish secular equilibrium. If the melt is moving faster than the solid matrix, then
U and Th will have different residence times within the
melting region because of their different bulk partition
coefficients. The less incompatible element (U) will spend
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Figure 4. Examples of how (230Th/238U), (226Ra/230Th),
and (231Pa/235U) activity ratios vary for a given set of
parameters for an equilibrium porous flow model in a midocean ridge setting in terms of porosity and mantle
upwelling velocity [from Spiegelman, 2000]. In principle,
it is possible to determine unique values for porosity and
upwelling velocity that are consistent with different
disequilibria measured in a single sample, but it is important
to realize that these results are model-dependent (both in
terms of the exact melting model used and in the choice of
partition coefficients).
more time in the melting column within the solid matrix
compared to the more incompatible element (Th) that is
preferentially partitioned into the faster moving melt. Further melting will continue to preferentially add the 230Th
ingrown from 238U in the source at a faster rate than 230Th
decay in the melt, because the melt travels faster than the
matrix, thus enhancing the 230Th-238U disequilibrium in the
final melt. As DTh > DRa and DU > DPa for mantle
compositions, then daughter ingrowth can also enhance
226
Ra-230Th and 231Pa-235U disequilibria in mantle melts.
The main difference between the various ingrowth melting
models is the extent to which the melt equilibrates with
the solid during melt extraction: the end-member models
being continuous equilibration [e.g., Spiegelman and
Elliott, 1993] and chemical isolation [e.g., McKenzie,
1985; Williams and Gill, 1989] (see section 5.1).
[28] As the different disequilibria pairs (230Th-238U,
226
Ra-230Th, and 231Pa-235U) respond differently to the
various parameters of melting models because of significant
differences in bulk partition coefficients and half-life, the
potential of the U series method will clearly be maximized
if all three pairs are measured on the same sample. Figure 4
shows an example [from Spiegelman, 2000] of how measurements of (230Th/238U), (226Ra/230Th), and (231Pa/235U)
from a single mid-ocean ridge basalt sample could potentially be used to constrain values for porosity and upwelling
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rate for a specific melting model (in this case the equilibrium porous flow model of Spiegelman and Elliott
[1993]). In general, the extents of 230 Th- 238 U and
231
Pa-235U disequilibrium are controlled mainly by the
melting rate, while the extent of 226Ra-230Th disequilibrium
is controlled mainly by the mantle porosity. The extent of
230
Th-238U disequilibrium is also strongly influenced by
the depth of melting due to differences in relative partitioning of U and Th in the different mineral assembleges
that are stable at different depths in the mantle. In the
shallow mantle (<1 GPa), clinopyroxene is the main host
phase for U and Th, and it has DTh > DU. However, at
higher pressures where aluminous clinopyroxene or garnet
are stable, the sense of U-Th fractionation changes, such
that Th is more incompatible than U (i.e., DTh < DU). This
contrasts with the behavior of 231Pa-235U disequilibria, as
Pa is thought to be more incompatible than U throughout
any mantle melting column.
[29] It is important to be aware that the absolute values
for mantle porosities and melting rates inferred from U
series disequilibria observations are critically dependent
both on the specific choice of assumed melting model as
well as on the values chosen for the bulk partition coefficients for Th, U, Pa, and Ra between melt and mantle. The
values for mineral/melt partition coefficients can vary as a
function of temperature and pressure, as well as mineral
composition, and so the bulk partition coefficient for an
element will vary both with the major element composition
of the mantle source and also with height within the melting
column.
2.4.4. Diffusion
[30] Most melting models assume that local chemical
equilibrium is maintained between the solid phases and
the melt, allowing trace element behavior to be described
through the use of equilibrium partition coefficients. During
melting the transfer of trace elements between solid phases
and melt will ultimately be controlled by solid-state diffusion. If the solid-state diffusion rate for a particular element
is very small, then there might not be sufficient time for
chemical equilibrium to be established prior to the melt
being extracted. Experimentally determined distribution
coefficients for U and Th in clinopyroxene [Van Orman et
al., 1998] indicate that diffusion of these large, highly
charged ions is very sluggish (1021 m2 s1 at 1200C),
such that chemical equilibrium might not be achieved
during melting. This will increase their effective partition
coefficients [e.g., Iwamori, 1993], such that they remain in
the solid phase longer during the melting process than is
predicted for simple chemical equilibrium. Several models
have been developed to account for the potential effects of
diffusion control on the U series nuclides during melting
[Qin, 1992, 1993; Iwamori, 1994; Van Orman et al., 1998,
2002].
[31] Our ability to assess the potential for differences
in solid-state diffusion rates to produce disequilibrium
between the U series nuclides is rather limited by the
general lack of experimentally determined distribution
coefficients for U, Th, Ra, and Pa in most minerals.
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observed in oceanic basalts might result from the diffusive
interaction of magmas percolating through older cumulates containing clinopyroxene (and/or plagioclase) within
the crust or at the crust-mantle transition zone. The faster
diffusion rate of 226Ra compared to 230Th and 238U
suggests that the 226Ra being produced in the crystal by
decay of 230Th could readily migrate out of the clinopyroxene crystal into the percolating melt, thus increasing
the (226Ra/230Th) of the melt.
Figure 5. U-Th equiline, or isochron, diagram illustrating
the kind of age information that can be obtained for
three different sets of samples. Samples that are older
than 5 times the 230Th half-life, i.e., 380 kyr, have
(238 U) = (230Th) and plot on the equiline with
(230Th/232Th) ratios that reflect their U/Th ratios (group 1).
Samples younger than 380 kyr can also plot on the
equiline provided that the processes that formed these
samples did not involve any fractionation of U and Th. If
samples plot off the equiline, then this indicates an event
that fractionated U and Th happened less than 380 kyr
ago. If samples define an isochron (group 2), their age can
be calculated as the slope equal to 1 elt, where l is the
230
Th decay constant and t is the age (note that the slope
of an isochron is constrained to values between 0 (t = 0)
and 1 (t = 1)). The assumptions are that at time t all these
samples had the same Th isotope ratio, and different U/Th
ratios, and that they would therefore have plotted on a
horizontal line. If samples plot to the right of the equiline,
they are said to have excess 238U, and if they plot to the
left, they have excess 230Th. Samples with excess 238U and
excess 230Th migrate vertically toward the equiline at a
rate that depends on the half-life of 230Th, and hence the
slope of the isochron equals 1 elt. The samples in
group 3 have variable (230Th/232Th) and almost constant
(238U/230Th), and so the time needed to move from the
sample with the greatest to the least (230Th/238U) can also
be calculated from the decay equation [e.g., Bourdon et
al., 2003b].
Van Orman et al. [2001] have developed a crystal lattice
elastic strain model that allows diffusion coefficients to be
estimated based on the ionic radius and charge of the
diffusing ion. This model predicts, for example, that Ra
should have a diffusion coefficient in clinopyroxene
(1017 m2 s1 at 1200C) that is 3 orders of magnitude
greater than that of either Th or U. This marked difference
has some intriguing implications for the development of
226
Ra-230Th disequilibrium in magmas [e.g., Saal and Van
Orman, 2004; Feineman and DePaolo, 2004]. Saal and
Van Orman [2004] have suggested that the 226Ra excesses
2.5. Extracting Time Information Graphically
From U Series Disequilibria
[32] For the 238U-230Th system, time information is
usually obtained with reference to the isochron, or equiline, diagram in which the parent (238U) and the daughter
(230Th) nuclides are normalized to an element or a longlived isotope (232Th) not involved in the decay scheme
(Figure 5), with all isotope ratios expressed as activity
ratios. The horizontal axis is equivalent to the elemental
U/Th ratio, and the vertical axis is the Th isotopic
composition, 230Th/232Th. All samples in secular equilibrium, where the activities of the parent and daughter
nuclides are equal, i.e., (238U)/(232Th) = (230Th)/(232Th),
will therefore plot on the 1:1 line (called the equiline) on
this isochron diagram [e.g., Allègre and Condomines,
1976]. Secular equilibrium can be disturbed by any chemical process that changes the relative abundances of the
parent and daughter nuclides, in this case U and Th. When
that process took place is determined by the difference
between the initial and the present-day Th isotope ratios.
The initial Th isotope ratio (230Th/232Th) is that of the
magma at the time of the event being studied, and it may be
significantly different from the Th isotope ratios of magmas
at the time of eruption due to subsequent radioactive decay
or ingrowth of 230Th. In contrast, other radiogenic isotope
systems such as 87Sr/86Sr would remain essentially invariant over such short timescales. The great strength of these
U series isotopes is that their isotope ratios change in
response to radioactive decay within the timescales of
magmatic processes, and age information is often obtained
in one of two ways.
[33] 1. If samples with different U/Th ratios plot on a
positive linear array (Figure 5), this array can potentially
represent an isochron, where the slope of the isochron will
correspond to an age (slope equal to 1 elt, where l is the
230
Th decay constant and t is the age, note that the slope of
an isochron is constrained to values between 0 (t = 0) and
1 (t = 1)). This age represents the time since the different
U/Th ratios were established, and it will be geologically
meaningful if the samples all had the same initial
(230Th/232Th) and remained as closed systems for U and
Th subsequently. The samples with different U/Th ratios
may be either suites of comagmatic whole rock lava
samples or mineral separates from an individual rock.
Mixing processes can also produce samples that fall on
a straight line, because the denominator (232Th) is the
same for both axes, but in this case the slope of the line
will not give a meaningful age. Thus it is important to
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establish that mixing between unrelated end-members has
not occurred. One potential indicator of mixing is significant compositional variations of long-lived radiogenic
isotopes such as 87Sr/86Sr that correlate with U/Th ratios.
[34] 2. Some comagmatic igneous rocks have similar
U/Th ratios but different Th isotope ratios, so that they
plot in near-vertical arrays on the U-Th equiline diagram
(Figure 5). In primitive rocks (i.e., that have experienced
minimal fractional crystallization) such vertical arrays may
reflect dynamic melting processes [e.g., McKenzie, 1985],
but in rocks related by fractional crystallization they may
reflect the time taken for magma differentiation to occur. If
we assume that the group 3 samples shown on Figure 5 are
related through different extents of fractional crystallization,
then the parental magma would have the high initial
(230Th/232Th) value, and the more evolved samples would
have lower (230Th/232Th) values because of radioactive
decay of 230Th as time elapses during the process of magma
differentiation.
[35] An analogous diagram to the U-Th equiline diagram
has been proposed for the 226Ra-230Th system using Ba as
the normalizing element as Ra does not have any long-lived
isotopes [e.g., Volpe and Hammond, 1991; Reagan et al.,
1992]. The Ba contents are essentially being used as a
means to estimate the amount of Ra initially present in
a melt or mineral. However, this diagram is more
difficult to interpret because the underlying assumption
that Ba is chemically similar to Ra is not strictly true
[e.g., Cooper et al., 2001]. The 5% difference in ionic
radius (Ra+ 1.42Å, Ba+ 1.48Å: eightfold coordination)
means that they will not partition identically into different
silicate minerals [e.g., Blundy and Wood, 2003], and
the implications of this are discussed in more detail in
section 4.3.2. For the 231Pa-235U system, Bourdon et al.
[1999b] have suggested using Nb as the normalizing
element for Pa. Despite their similar ionic charge and
broadly similar geochemical behavior the difference in
ionic radius (Pa5+ 0.78Å, Nb5+ 0.64Å: sixfold coordination) will also lead to different crystal partitioning behavior
[Blundy and Wood, 2003].
3. U SERIES DISEQUILIBRIA IN IGNEOUS ROCKS
FROM DIFFERENT TECTONIC SETTINGS
3.1. Data Compilation
[36] There are now sufficient high-quality data available
in the literature to do an integrated review of global
230
Th-238U, 226Ra-230Th, and 231Pa-235U disequilibria in
igneous rocks. This allows us to establish the observed
range in values for these disequilibria in nature and also to
highlight differences in disequilibria found in magmas
generated in different tectonic environments. The results
of this new compilation are summarized in Figures 6 and 7.
We have restricted the data to mass spectrometer analyses,
except for some of the 226Ra data, and we have concentrated on analyses of samples that appear to have been little
affected by alteration and shallow level processes. For
simplicity, samples are divided into the three tectonic
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environments that represent the principal melt generation
regimes on the planet: (1) mid-ocean ridges, (2) withinplate settings (ocean island basalts and continental intraplate lavas), and (3) subduction zones. Figure 6 shows a
series of histograms of (230Th/238U), (226Ra/230Th), and
(231Pa/235U) activity ratios for volcanic rocks from different
tectonic environments (data sources are given in Figure 6
caption). Figure 7a shows the U-Th equiline diagram (a plot
of the activities of 230Th versus 238U, both normalized to
the activity of the long-lived 232Th isotope), which highlights the extent of element U-Th fractionations. Figures 7b
and 7c show the covariations between (230Th/238U) disequilibria and (231Pa/235U) and (226Ra/230Th) disequilibria,
respectively.
[37] All data are plotted as initial ratios calculated back to
the eruption age, except for the submarine MORB and back
arc basin samples for which the eruption ages are generally
unknown. The short t1/2 of 226Ra (1.6 kyr) means that it is
reasonable to assume that any MORB sample with
226
Ra-230Th disequilibrium is younger than 8000 years
and that any correction for posteruptive decay of 230Th or
231
Pa would be negligible. Those MORB samples with
226
Ra- 230 Th disequilibrium are distinguished on the
230
( Th/238U) and (231Pa/235U) histograms in Figure 6. Furthermore, for the MORB samples plotted on Figures 6 and
7, the (226Ra/230Th) values only represent minimum estimates for the (226Ra/230Th) disequilibria on eruption because of the unknown extent of posteruptive decay (except
for a few samples of known age collected by submersible
from the East Pacific Rise [Sims et al., 2002a]).
[38] The most extreme disequilibria values of any igneous rocks are found in historic natro-carbonatite samples
from Oldoinyo Lengai volcano, East Africa [Williams et al.,
1986; Pyle et al., 1991; Pickett and Murrell, 1997]. These
rare carbonate-rich melts have (230Th/238U) of 0.1 – 0.2,
(226Ra/230Th) of 40– 80, and (231Pa/235U) of 0.18– 0.22.
These data are consistent with models in which the natrocarbonatite magma forms by immiscibility from a silicate
nephelinite magma, where Ra and U are strongly partitioned
preferentially into the natro-carbonatite magma relative to
Th and Pa. These lavas are further unique in being the only
lava samples where 228Ra-232Th disequilibrium has been
convincingly measured, with (228Ra/232Th) of 27. Given
the short half-life of 228Ra (5.75 years), such large disequilibria indicate a very short time between Ra-Th fractionation
and lava eruption: 7 – 18 years for the 1960 – 1963 A.D.
lavas [Williams et al., 1986] and 20– 80 years for the 1988
A.D. lavas [Pyle et al., 1991].
3.2. The 230Th-238U Disequilibria
[39] Mid-ocean ridges and within-plate settings are both
characterized predominantly by lavas with 230Th excesses
(Figure 6). Within-plate lavas show a greater range in
(230Th/238U) than MORB lavas, with values generally
between 1.0 and 1.6, although a few basalts from the
western United States do have slight 238U excesses [Reid
and Ramos, 1996]. Most MORB lavas have (230Th/238U)
values between 1.0 and 1.3. MORB samples with
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Figure 6. Histograms of measured (230Th/238U), (226Ra/230Th), and (231Pa/235U) disequilibria in
volcanic rocks from different tectonic settings: mid-ocean ridge basalts (MORB); within-plate, ocean
island basalts and continental within-plate lavas; and ARCS, subduction zone lavas. Dotted line indicates
secular equilibrium, i.e., samples with (230Th) = (238U), (226Ra) = (230Th), or (231Pa) = (235U). All data are
obtained by thermal or plasma ionization mass spectrometry, except some (226Ra/230Th) data. Data
sources are McDermott and Hawkesworth [1991], Goldstein et al. [1992], Williams et al. [1992], Cohen
and O’Nions [1993], Goldstein et al. [1993], Volpe and Goldstein [1993], Reagan et al. [1994], Asmerom
and Edwards [1995], Lundstrom et al. [1995], Reid [1995], Bourdon et al. [1996a, 1996b], Cohen et al.
[1996], Reid and Ramos [1996], Sigmarsson [1996], Turner et al. [1996], Elliott et al. [1997], Huang et
al. [1997], Pickett and Murrell [1997], Regelous et al. [1997], Turner et al. [1997a, 1997b], Widom et al.
[1997], Bourdon et al. [1998], Clark et al. [1998], Claude-Ivanaj et al. [1998], Lundstrom et al. [1998b],
Sigmarsson et al. [1998a], Turner et al. [1998], Asmerom [1999], Chabaux et al. [1999], Lundstrom et al.
[1999], Sims et al. [1999], Thomas et al. [1999], Turner et al. [1999], Vigier et al. [1999], Bourdon et al.
[2000a, 2000b], Sturm et al. [2000], Turner et al. [2000a, 2000b], Claude-Ivanaj et al. [2001], Cooper et
al. [2001], Peate et al. [2001a, 2001b], Pietruszka et al. [2001], Turner et al. [2001a], Turner and Foden
[2001], Cooper et al. [2002], Sims et al. [2002a], Thomas et al. [2002], Cooper et al. [2003], Dosseto et
al. [2003], Fretzdorff et al. [2003], George et al. [2003], Kokfelt et al. [2003], Lundstrom et al. [2003],
Stracke et al. [2003], Turner et al. [2003b], Yokoyama et al. [2003], Zou et al. [2003], and Tepley et al.
[2004]. Samples that appear to have been affected by alteration or shallow level contamination processes
have been filtered out. All data are plotted as initial ratios calculated back to the eruption age, except for
submarine mid-ocean ridge and back arc basin lavas for which eruption ages are generally unknown. For
these settings, samples with measured 226Ra-230Th disequilibrium are distinguished on the (230Th/238U)
and (231Pa/235U) histograms as black. These samples are inferred to have been erupted within the last
8000 years, and thus there will have been minimal radioactive decay of 230Th and 231Pa since eruption.
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(230Th/238U) < 1 are rare, but a few examples are known from
the Mid-Atlantic Ridge at 29– 30N [Bourdon et al.,
1996b], the Kolbeinsey Ridge [Sims et al., 2002b], and the
Garrett Transform [Tepley et al., 2004]. In contrast, many
subduction zone lavas have (230Th/238U) < 1 [e.g., Gill and
Williams, 1990], with values as low as 0.4, although about a
quarter of the samples compiled for Figures 6 and 7 have
(230Th/238U) > 1 similar to MORB and within-plate lavas.
The horizontal axis on Figure 7a is (238U/232Th), which is
directly proportional to the elemental U/Thatomic ratio in the
sample. Subduction zone lavas clearly show the widest
range in (238U/232Th) from 0.4 to 3.4, with the highest
values generally found in the more trace element-depleted
(i.e., lower Th content) samples [e.g., McDermott and
Hawkesworth, 1991]. MORB lavas have (238U/232Th)
between 0.9 and 1.6. Within-plate lavas generally have
lower (238U/232Th) values than MORB, between 0.4 and
1.2, with the lowest values (0.4– 0.7) found in highly
enriched continental potassic lavas (Gaussberg, Antarctica
[Williams et al., 1992]; Nyamuragira, East Africa [Pickett
and Murrell, 1997]; Wudalianchi, China [Zou et al.,
2003]; and Tibet [Cooper et al., 2002]).
3.3. The 231Pa-235U Disequilibria
[40] With the exception of some depleted arc tholeiite
lavas from the Tonga-Kermadec island arc [Bourdon et al.,
1999b] and the Oldoinyo Lengai carbonatite lavas [Pickett
and Murrell, 1997] all lava samples show 231Pa excesses
irrespective of their tectonic setting. MORB lavas generally
have the greatest observed excesses, with (231Pa/235U)
values from 1.9 to 4.0. Within-plate lavas, including both
ocean island basalts and continental-intraplate lavas, usually
have lower (231Pa/235U) values from 1.1 to 2.4. Subduction
zone lavas have even lower (231Pa/235U) values from 0.8 to
1.7, with the lowest values found in lavas with the lowest
(230Th/238U) values (Figure 7b). Kick ’em Jenny volcano in
the Lesser Antilles is an exception, with the highest
(231Pa/235U) value of 2.15 and yet one of the lowest
(230Th/238U) values of 0.6 [Pickett and Murrell, 1997].
Subduction zone lavas and within-plate lavas form broad
positive arrays on Figure 7b, with MORB lavas displaced to
higher (231Pa/235U).
Figure 7. Global variations in U series disequilibria
in igneous rocks from different tectonic settings:
(a) (238U/232Th) versus (230Th/232Th), (b) (230Th/238U)
versus ( 23 1 Pa/ 2 35 U), and (c) ( 23 0 Th/ 2 3 8 U) versus
(226Ra/230Th). Data sources are as for Figure 6. Black
dotted lines on Figure 7a indicate different extents of
230
Th and 238U excesses in percent: For example, 50%
230
Th excess (to the left of the equiline) indicates
(230Th/238U) of 1.5, whereas 50% 238U excess (to the
right of the equiline) indicates (238U/230Th) of 1.5. Shaded
dotted lines indicate secular equilibrium.
3.4. The 226Ra-230Th Disequilibria
[41] In general, mafic igneous rocks with (226Ra/230Th) <
1 are very rare. Subduction zone lavas show the greatest
range in (226Ra/230Th) values from 1 to almost 7, with
the highest values mainly associated with the lowest
(230Th/238U) values (Figure 7c). MORB lavas mostly
have (226Ra/230Th) values between 1 and 3, with a few
exceptional samples having values up to 4.2. Withinplate mafic lavas have a relatively restricted range in
(226Ra/230Th), with values from 0.8 to 1.6, and it is only a
few oceanic island basalts (OIB) samples from the Azores
and Samoa [Claude-Ivanaj et al., 2001; Bourdon and Sims,
2003] that have (226Ra/230Th) < 1. Evolved lavas (e.g.,
phonolites from Mount Erebus, Antarctica [Reagan et al.,
1992]; trachytes from Azores [Widom et al., 1992]; and
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dacites from Santorini, Greece [Zellmer et al., 2000]) often
have significant 226Ra deficits, e.g., (226Ra/230Th) 0.75 –
0.92, as a result of extensive fractional crystallization of
feldspar, which preferentially removes Ra from the melt
relative to Th.
[42] As most magmas have experienced some degree of
differentiation, we look first at the timescales of magma
differentiation and crystallization within the crust, which are
potentially linked to the eruptive behavior of volcanoes.
Next we address melt generation processes at mid-ocean
ridges and ocean islands and then finish with the processes
of fluid transfer from the downgoing slab and melt generation in the more complex setting of subduction-related
magmatism.
grains, and it is difficult to determine the textural relations
of the individual crystals. Exceptions include accessory
minerals with high-U and/or high-Th contents such as
zircon and allanite for which in situ 238U-230Th ages can
be determined by ion microprobe [e.g., Reid et al., 1997;
Charlier et al., 2003; Vazquez and Reid, 2004]. The
general principles of how U series disequilibria are used
to establish timescales of magma chamber processes and to
date young volcanic rocks are reviewed by Condomines et
al. [2003] and Reid [2003]. Hawkesworth et al. [2000,
2004] provide broader overviews that integrate the U
series results with other methods of determining the timescales of magmatic processes and then discuss the wider
implications of such information.
4. TIMESCALES OF MAGMATIC DIFFERENTIATION
AND CRYSTAL GROWTH
4.2. Magmatic Differentiation Rates and Crustal
Residence Times From Whole Rock Samples
4.2.1. Magma Residence Times
and Crustal Transit Times
[45] Early attempts to infer magma residence times from
U series nuclides developed models for steady state magma
chamber systems, in which the eruptive volcanic output
effectively balances the influx of new magma batches so
that the chamber volume remains essentially constant.
These models have been applied to persistently active
centers such as Hawaii, Etna, Stromboli, and Reunion
[Pyle, 1992; Albarède, 1993; Condomines, 1994; Hughes
and Hawkesworth, 1999; Condomines et al., 2003]. The
extent of isotopic disequilibrium in the erupted lavas from
a steady state reservoir will be roughly constant and
governed by the balance between higher values due to
input of fresh magma to the chamber and lower values due
to radioactive decay during residence of the mixed magma
in the chamber prior to eruption. Provided that the initial
degree of isotopic disequilibrium in the input magma is
known, often from some flank eruptive unit that bypassed
the chamber, an average magma residence time can be
determined. Residence times of tens to hundreds of years
are typically obtained, with 10% of the volume of the
magma chamber inferred to be erupted annually. Such
magma residence times may or may not be accompanied
by significant magma differentiation, and so different
approaches are required to distinguish rates of magma
differentiation from the time taken for magma to traverse
the crust.
[46] If the degree of isotopic disequilibrium of a magma
when it leaves the melting zone can be inferred, then its
isotope composition at the time of eruption can be used to
constrain how long it took to travel through the crust. The
simplest case is for magmas erupted with significant 226Ra
excesses, as these imply crustal residence times of less than
8000 years, provided that the 226Ra-230Th disequilibrium
was established by mantle melting processes. The same
arguments can be made for samples with 230Th-238U disequilibria, but the longer half-life of 230Th means that
it constrains the crustal residence ages to be less than
380,000 years. This provides much less insight given the
much shorter timescales for variability within the plumbing
4.1. Background
[43] Differentiation of mantle-derived magmas occurs
principally by fractional crystallization, where early formed
crystals are removed from the magma, thereby changing its
composition [Bowen, 1928]. The time required for such
crystal-liquid separation to occur will depend on the density
contrast between the crystals and the liquid, the viscosity of
the liquid, the size of the crystals, and the dynamics of the
magma chamber system. The causes and timescales of
crystallization vary: (1) Crystallization in response to cooling will depend on the size of the magma body, the magma
replenishment rate, and the thermal structure of the crust
[e.g., Annen and Sparks, 2002]. (2) Crystallization in
response to degassing and decompression may be fast (days
or weeks), and there is increasing evidence that it is too fast,
and occurs too close to the time of crystallization of the host
rock, for crystals formed in this way to be involved in
fractional crystallization [e.g., Zellmer et al., 2003]. Thus
magma differentiation processes may be largely thermally
controlled. Progressive crystallization will result in the
buildup of volatiles in the melt, and the sudden release of
these dissolved volatiles may, in turn, trigger explosive
eruptions. Thus there is considerable interest in the links
between crystallization and the timing and style of volcanic
eruptions and also in the primary controls on magma
differentiation.
[44] U series disequilibria are just one way to determine
the timescales of magmatic processes; others include
relative chronometers such as crystal size distributions
and compositional profiles across crystals that may have
been modified by diffusion [e.g., Cashman, 1990;
Davidson and Tepley, 1997; Zellmer et al., 1999;
Morgan et al., 2004]. The latter provide information
on how long crystals have been at magmatic temperatures
rather than when that might have been, and age information is obtained on individual crystals so that the age
profiles of crystal populations can be determined for
individual rocks. U series disequilibria yield absolute ages,
but in most cases they are determined on mineral separates
that may comprise several tens to hundreds of crystal
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226
Figure 8. Variations of (226Ra/230Th) with Th content (in
ppm), an index of differentiation, in three suites of
comagmatic lavas: (1) Asal rift, Afar [Vigier et al.,
1999]; (2) Surtsey-Heimaey, Iceland [Sigmarsson, 1996];
(3) Miyakejima volcano, Japan [Yokoyama et al., 2003].
(230Th/238U) and (87Sr/86Sr) are both constant within the
Afar and Iceland suites, indicating that the samples within
each suite are simply related by different degrees of
differentiation from similar parental magmas. The slope
of the arrays will therefore depend on the rate of
differentiation relative to the decay of 226Ra and on the
degree of Ra-Th fractionation during fractional crystallization. However, (230Th/238U) is not constant for the
historic samples from Miyakejima but increases from
0.755 to 0.793 as (226Ra/230Th) decreases, and this is
consistent with control by magma mixing rather than
simple closed system differentiation.
systems beneath many volcanoes inferred from stratigraphic
records.
4.2.2. Estimates of Magma Differentiation Rates
From (226Ra/230Th)
[47] A widely used approach to determine rates of
magma differentiation is to investigate how the magnitude
of U series disequilibria, from which age information can
be obtained, varies with indices of differentiation [e.g.,
Hawkesworth et al., 2000]. Figure 8 shows (226Ra/230Th)
activity ratios plotted against Th contents as an index of
differentiation for three magmatic suites (Asal rift, SurtseyHeimaey, and Miyakejima). As Th is normally an incompatible element during crystallization, then Th contents
will increase in the magma as differentiation proceeds and
progressively more crystals are removed from the magma.
The samples all have (226Ra/230Th) > 1 and will evolve to
secular equilibrium, i.e., (226Ra/230Th) = 1, by radioactive
decay. (226Ra/230Th), and hence the degree of 226Ra-230Th
isotope disequilibrium, decreases with increasing Th
contents in these volcanic suites, suggesting that the
more evolved magmas are ‘‘older’’ than the less evolved
magmas. However, interpretations are complicated by the
fact that Ra/Th ratios in magmas can potentially be changed
by progressive removal of feldspar crystals during magma
differentiation. The slope of the arrays will therefore depend
both on the rate of differentiation relative to the decay of
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Ra and on the relative bulk partition coefficients for Ra
and Th during fractional crystallization. The latter will be
governed by the modal proportion and composition of
feldspar in the assemblage of crystals being removed from
the magma.
[48] 1. For the Asal rift lavas the 30% fractional crystallization of a plagioclase-dominated assemblage necessary to
explain the major and trace element variations cannot
explain all of the observed range in (226Ra/230Th), which
therefore must also reflect different magma residence times,
even though the samples were all erupted within a one week
period in 1978 A.D. [Vigier et al., 1999]. Several magma
batches must have been injected into one or more reservoirs
beneath the rift at different times prior to eruption and then
differentiated for between 1000 and 2000 years, depending
on whether they evolved as independent closed systems or
as an open system zoned magma chamber replenished by
each new batch of primitive magma.
[49] 2. In contrast, Sigmarsson [1996] argued that plagioclase fractionation alone could explain the differences
in (226Ra/230Th) between the 1963 and 1967 A.D. Surtsey
alkali basalts and the 1973 A.D. Heimaey hawaiitemugearite lavas in the Vestmannaeyjar volcanic system,
southern Iceland, and variations of the short-lived 210Pb
nuclide (t1/2 22 years) instead indicated rapid differentiation from alkali basalt to hawaiites and mugearites in only
10 years or so. He proposed a model in which a small
volume of alkali basalt magma was injected into a deep
(15 – 20 km) reservoir in relatively cold crust beneath
Heimaey at the same time as similar magmas were erupted
at Surtsey. The relatively cold crust led to rapid cooling
and differentiation to hawaiites and mugearites, and the
consequent buildup of volatiles led to the 1973 A.D.
eruption.
[50] 3. For the historic lavas from Miyakejima volcano, Japan [Yokoyama et al., 2003], the variations in
(226Ra/230Th) correlate with (230Th/238U) and thus reflect
magma mixing processes rather than timescales of
differentiation.
4.2.3. Estimates of Magma Differentiation Rates
From (238U/230Th)
[51] The longer-lived 238U-230Th system has been used to
determine the timescales of closed system differentiation
from mafic parental magmas to felsic evolved magmas for
several different magmatic suites [e.g., Reagan et al., 1992;
Widom et al., 1992; Bourdon et al., 1994; Bohrson and
Reid, 1998; Hawkesworth et al., 2000]. In each case it is
assumed that the 230Th-238U disequilibrium measured in an
erupted mafic sample is the same as that of the mafic
magma prior to its differentiation into a more evolved
magma. The difference in initial (230Th/232Th) between
the mafic magma and a differentiated magma with similar
U/Th in a closed system simply represents radioactive
decay, thus allowing the duration of the magma differentiation to be calculated. For example, Bourdon et al. [1994]
estimated a differentiation time of 100 kyr for the Laacher
See phonolite (Germany) to evolve from a possible parental
basanite magma (Figure 9), and Widom et al. [1992]
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Figure 9. Example of different timescales of magma
differentiation, the circa 13 ka Laacher See zoned phonolite
eruption, Eifel, Germany [Bourdon et al., 1994]. These data
are consistent with a two-stage model in which differentiation of a parental basanitic magma in a deep crustal
reservoir takes 100 kyr and then the resulting mafic
phonolite magma is emplaced into a shallow level magma
chamber where it differentiates to more evolved compositions over a period of <20 kyr. The composition of the
initial parental basanitic magma (open circle) was estimated
from the (230Th/232Th) ratio of an olivine xenocryst in the
phonolite and the (238U/232Th) of a local older basanite lava.
The difference in (230Th/232Th) between the parental
basanite and the mafic phonolite can be explained by
radioactive decay during 100 kyr of magma differentiation, with no change in U/Th of the magma. Subsequent
differentiation within the shallow magma chamber leads to
increasing (238U/232Th) in phonolitic pumice glass samples
(shaded circles) because of progressive removal of a mineral
assemblage containing accessory phases, apatite, and
sphene, with high Th and U contents and low U/Th ratios
(see Figure 10b). These phonolitic glasses define an
isochron with an age of 14.3 ± 6.5 ka that is within error
of the known eruption age as determined from 238U-230Th
mineral isochrons, 14C dating, and 40Ar/39Ar dating.
Figure 9 is modified from Bourdon et al. [1994], with
permission from Elsevier.
inferred that it took 80 kyr to produce trachytic magmas
from a parental alkali basalt at Agua de Pão, Sao Miguel, in
the Azores.
[52] The basanite to phonolite sequence on Tenerife
(Canary Islands) can be modeled by 80% largely closed
system fractional crystallization, but this differentiation
occurred in distinct stages at different crustal levels: Mafic
and intermediate lavas fractionated in a deep crustal reservoir at 6 – 9 kbar and then subsequently differentiated to
evolved phonolitic compositions in a shallow chamber at
1.5 – 2.5 kbar [Ablay et al., 1998]. Variations of initial
(230Th/238U) values with Zr, an index of differentiation,
suggest that the extensive fractionation in the deep crust
took much longer than that at shallow levels [Thomas,
1998; Hawkesworth et al., 2000]. While the decrease in
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(230Th/238U) with differentiation indicates an overall
differentiation time of 230 kyr, whole rock 226Ra-230Th
data suggest that the final feldspar-dominated differentiation
to the most evolved phonolites took place within a thousand
years or so before eruption. However, Lundstrom et al.
[2003] measured significant 226Ra excesses in three lavas
from Tenerife that spanned a large compositional range from
9 to 2 wt % MgO. This would require a much shorter
differentiation time than suggested by the 230Th-238U
system, perhaps implying that the 230Th-238U disequilibria
were affected by open system processes. One difficulty
with this is that these lavas show a striking increase and
then decrease in Ba with increasing differentiation, as
reflected in Zr contents, and thus there seems to have
been no simple mixing between high- and low-MgO
magmas that might have a crustal origin.
[53] Assimilation of basement rocks during differentiation in the crust is always a possibility in both continental
and oceanic environments, and it will tend to give lower
238
U-230Th disequilibrium in the evolved magma as the
assimilated material is likely to be old enough to be in
secular equilibrium and lie on the equiline. Therefore, in
such open systems the timescales for differentiation will be
overestimated because the reduction in measured disequilibrium in the evolved magmas is due to the combination of
radioactive decay and addition of material in secular equilibrium. Bohrson and Reid [1998] showed that rhyolites on
Socorro Island, Mexico, are related to trachytes by fractional crystallization, but systematic variations in 87Sr/86Sr
indicated that differentiation was accompanied by assimilation of seawater-altered silicic basement rocks, and thus
the 40 kyr differentiation time indicated by the difference in (230Th/232Th) is a maximum estimate. If the
assimilants are relatively young plutons or crystal mushes,
perhaps related to an earlier phase of the same overall
magmatic system, then this assimilation might be difficult
to recognize with conventional geochemical monitors such
as Sr and Pb isotopes because not enough time has
elapsed for the isotope compositions to respond to
changes in Rb/Sr, U/Pb, and Th/Pb ratios. However, if
this crust has a markedly different U/Th value, it will
develop a distinctive (230Th/232Th) value over just a few
hundred thousand years. This ‘‘cryptic’’ assimilation has
been proposed to explain the 232Th-230Th-238U variations in
some continental arc magmas from the Cascades and the
Aleutians [George et al., 2003; Reagan et al., 2003].
[54] It is important to bear in mind that not all silicic
magmas are formed by extensive fractional crystallization
from a mafic parental magma. They can also be produced
by partial melting of crustal material, and U series disequilibria can sometimes help to differentiate between these two
mechanisms [e.g., Sigmarsson et al., 1991]. For example,
the trachytes on Socorro Island (Mexico) cannot be related
to contemporaneous basaltic magmas through closed system
fractional crystallization as they, in fact, have greater
230
Th-238U disequilibrium at a given (238U/232Th) than the
basalts. Bohrson and Reid [1998] proposed a model to
generate the trachytes by partial melting of basaltic base-
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ment that could explain both the U series disequilibria data
as well as the major and trace element variations.
4.2.4. Rapid Magma Differentiation Timescales
in Evolved Magmatic Systems
[55] Shorter differentiation timescales are often associated with the development of zonation in silicic magma
chambers. About 80% fractional crystallization is necessary to explain the compositional range of phonolitic
glasses from the circa 13 ka Laacher See zoned eruption,
and yet they define a 238U-230Th isochron on Figure 9
with an age of 14.3 ± 6.5 ka. This age is within error of
the eruption age and therefore requires that the differentiation within the phonolitic magma occurred rapidly and
shortly before eruption. Detailed modeling gives a maximum timescale of 10– 20 kyr for the development of this
compositional zonation, consistent with thermal and fluid
dynamical models for the Laacher See magma chamber
[Bourdon et al., 1994]. The extreme chemical zonation
preserved in trachytic eruptive units from Agua de Pão
in the Azores requires 70% fractional crystallization
dominated by alkali feldspar, and Widom et al. [1992]
estimated that this zonation took less than 4600 years to
develop, based on (226Ra/230Th) data and stratigraphic
constraints.
[56] Rogers et al. [2004] analyzed 238U-230Th-226Ra isotope ratios in a stratigraphic sequence of aphyric trachytic
lava flows from Longonot volcano, Kenya, that were erupted
in the last 6000 years. These lavas record progressive closed
system fractional crystallization, dominated by alkali feldspar in which Ra and Ba are highly compatible. The
(230Th/238U) disequilibria indicate that fractionation of
U/Th took place within the last 10,000 years, as the
samples all fall on a broadly horizontal trend (i.e.,
constant (230Th/232Th)) on an equiline diagram. Rogers et
al. [2004] developed a model that combined radioactive
decay and fractional crystallization to estimate rates of
differentiation from the (226Ra/230Th) data. Differentiation
rates for evolution from hawaiite to trachyte were about
0.2 104 year1, with faster rates for differentiation
within the trachytes of approximately 3 104 year1.
The (226Ra/230Th) data on whole rock and alkali feldspar
separates indicated that phenocryst formation continued
almost up to the time of eruption.
4.3. Timing of Crystallization From Mineral Separates
4.3.1. Discordance Between Time of Crystal Growth
and Time of Lava Eruption
[57] The incentive behind early studies of 238U-230Th
disequilibria in mineral separates was as a means of determining the eruption age of young volcanic rocks using an
isochron approach: Different mineral phases with a range of
U/Th ratios, separated from an individual sample, should
define an isochron age on the equiline diagram (Figure 5).
Results were mixed, as 238U-230Th isochron ages were often
much older than independently known eruption ages. It was
soon realized that such discordant data actually provided
useful information about the timing of crystal growth, which
did not necessarily coincide with the eruption age.
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[58] Analysis of fine-grained groundmass crystals that
grew after eruption during rapid surface cooling of a lava
flow should provide an unequivocal determination of the
eruption age, and Peate et al. [1996] determined a
230
Th-238U isochron age of 156±29 ka on groundmass
magnetite separates (grain size of 20 mm) and whole rock
samples from a sequence of lavas from Albuquerque, New
Mexico, that record a Quaternary geomagnetic excursion
(Figure 10a). However, most studies have analyzed phenocryst phases that likely began to grow within the magma
chamber or plumbing system prior to lava eruption.
[59] For samples where 230Th-238U dating of phenocrysts
gives the same age as independently determined eruption
ages, this implies that there was a very short time (relative to
the 230Th half-life) between crystal growth and eruption.
For the Laacher See zoned eruption, three pumice samples
from the crystal-rich mafic phonolitic units give internal
230
Th-238U mineral isochrons with a mean age of 13 ±
3 ka (Figure 10b), identical to the eruption age, indicating
a maximum residence time for the crystals in the magma
chamber of 1 – 2 kyr [Bourdon et al., 1994].
4.3.2. Timescales of Crystal Residence and Growth
From 226Ra/230Th Disequilibria
[60] For those cases where 230Th-238U mineral isochrons
give ages within error of the known eruption age (e.g.,
Mount St. Helens andesite (Figure 11a) [Volpe and
Hammond, 1991]), this implies that crystal growth must
have happened less than a few thousand years before
eruption. As these timescales are of a similar magnitude to
the 226Ra half-life (1599 years), the 226Ra-230Th system
can potentially provide better estimates of the crystal
residence and growth times if the samples are young
enough. Obtaining age information using the 226Ra-230Th
system on separated components of a rock is not as easy
as for the 230Th-238U system, because interpretation of
data on a Ba-normalized 226Ra-230Th isochron plot, as
alluded to earlier, is not straightforward. Volpe and
Hammond [1991] showed that for andesite samples from
Mount St. Helens the groundmass and whole rock analyses
always plotted above a reference line through the mineral
phases (Figure 11b), which suggested that some open
system behavior had taken place. However, the main
assumption behind the construction of a diagram like
Figure 11b, that Ba and Ra have identical geochemical
behavior, is probably incorrect [e.g., Cooper et al., 2001;
Blundy and Wood, 2003].
[61] Cooper et al. [2001] devised an elegant method to
determine crystal residence ages from 226Ra-230Th-Ba
data that takes into account the different mineral/melt
partitioning of Ba and Ra. They used an elastic strain
partitioning model [Blundy and Wood, 2003] to estimate
partition coefficients for Ba and Ra in different mineral
phases under conditions relevant to the magma under
investigation. Differences in partition coefficients mean that
the mineral phases will have different initial (226Ra)/Ba
ratios. They plotted (226Ra)/Ba versus time evolution curves
for the groundmass and for melts in equilibrium with each
analyzed crystal phase (using the estimated partition
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coefficients). If the curves intersect, then this defines the
time when the minerals could have crystallized from the
host magma. It is important to correct for the effects of
impurities in the analyzed bulk mineral separates (adhering glass and melt inclusions) by mass balance calculations using the difference between Th and Ba contents in
the bulk mineral separates and those in the pure minerals
as determined by in situ ion probe measurements. For the
Mount St. Helens andesite example shown in Figure 11c,
this approach gives a residence time of 2.4 –4.4 ka for
plagioclase and 1.5– 5.7 ka for pyroxene prior to eruption
[Cooper and Reid, 2003].
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[62] Cooper et al. [2001] used this technique to show
that 226Ra-230Th-Ba data from the 1955 A.D. east rift
eruption at Kilauea volcano (Hawaii) are consistent with
cocrystallization of plagioclase and pyroxene from a melt
represented by the groundmass with a mean age of 1000 ±
400 years. Another example is provided by Zellmer et
al. [2000], who studied the Kameni dacites (46 A.D. to
1950 A.D.) on Santorini (Aegean arc) and determined an
essentially zero age 230Th-238U mineral isochron (18 ±
18 ka) for a sample of the 1950 A.D. lava. Modeling of
226
Ra-230Th-Ba data in whole rock Kameni dacites
indicated that plagioclase fractionation occurred less than
1000 years prior to each individual eruption, which is
consistent with the short (maximum of a few hundred
years) plagioclase crystal residence times calculated from
Sr diffusion profiles [Zellmer et al., 1999] and crystal
size distribution studies [Higgins, 1996].
4.3.3. Interpretations of Discordant Ages
[63] There are numerous examples in the literature where
230
Th-238U mineral isochrons give well-defined ages that
are older than the known eruption age: (1) Volpe [1992]
obtained ages of 27 ± 18 ka and 28 ± 10 ka from two
Hotlum andesite samples from Mount Shasta (Cascades,
United States) that were erupted less than 4000 years ago;
(2) Volpe and Hammond [1991] measured old ages of 34 ±
16 ka in a basalt lava and 27 ± 12 ka in an andesite lava
from the Castle Creek eruptive period (1.7 – 2.2 ka) at
Mount St. Helens (Cascades, United States); and (3) Heath
et al. [1998] analyzed four separate lava samples younger
than 4000 years old from Soufriere volcano (St. Vincent,
Lesser Antilles) that all yielded old mineral isochron ages of
between 46 and 77 ka.
[64] Minerals with old ages may be xenocrysts incorporated by the magma during ascent to the surface and as such
Figure 10. (a) The 230Th-238U isochron defined by whole
rock and groundmass magnetite (20 mm crystals) from a
suite of related basaltic lavas from the Albuquerque
volcanic field, New Mexico, United States, that records an
eruption age of 156 ± 29 ka [Peate et al., 1996]. (b) Internal
mineral isochrons for pumice samples from the main part
of the zoned Laacher See phonolite eruption (solid circles)
indicating a crystallization age of 13 ± 3 ka, indistinguishable from the eruption age and the isochron defined by
glass samples (see Figure 9). Separates from the most
evolved, crystal-poor samples give an older and less
precise age of 30 ka, suggesting incorporation of crystals
from older cumulates [Bourdon et al., 1994]. (c) Pyroxene
and magnetite separates from an 190 ka basanite lava
flow (Fornicher Kopf, Eifel, Germany) lying on the
equiline, whereas the groundmass has 5% excess 230Th.
Tie lines between either mineral phase and the groundmass
have slopes greater than the equiline and thus give
indeterminate ages. This implies that these mineral phases
are old (>380 ka) and not in equilibrium with the host
groundmass and are likely to be xenocrysts perhaps
inherited from an earlier magmatic episode [Peate et al.,
2001c]. Abbreviations are gnd, groundmass; gl, glass; wr,
whole rock; mt, magnetite; pyx, pyroxene; ap, apatite; sp,
sphene; am, amphibole; and ha, haüyne.
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Figure 11. Crystallization ages of a Mount St. Helens
andesite (0.4 ka, Cascades, United States). (a) The
230
Th-238U equiline diagram (modified from Volpe and
Hammond [1991], with permission from Elsevier). (b) Banormalized 226Ra-230Th isochron diagram (modified from
Volpe and Hammond [1991], with permission from Elsevier).
The whole rock does not lie on the pl-pyx tie line, which
suggested open system behavior to Volpe and Hammond
[1991], but it is more likely a consequence of the different
partitioning of Ra and Ba between these minerals and melt
[Cooper and Reid, 2003]. (c) The 226Ra/Ba time evolution
diagram [Cooper and Reid, 2003], showing that the
plagioclase crystallized from the host melt at 3 ka (shaded
area). However, the timing of pyroxene growth is less well
constrained (1.5 – 5.7 ka) (modified from Cooper and Reid
[2003], with permission from Elsevier).
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provide little insight into magma residence times or the
timescales of differentiation. Alternatively, they may be
minerals that crystallized from earlier magma batches and
so constrain the length of time over which a particular
magma system has been active. It is possible that the old
ages reflect a significant residence time of the crystals in the
magma prior to eruption. However, at least for the examples
given above, closed system residence times >10 kyr can be
ruled out because the magmas all show 226Ra-230Th
disequilibrium that would have returned to secular equilibrium after only 8 kyr, and thus some type of open
system behavior is required.
[65] Analyses of different mineral phases do not always
lie on simple linear trends on the 230Th-238U isochron
equiline diagram, suggesting that each phase might have
crystallized at different times. For example, Peate et al.
[2001c] analyzed several different phases separated from
two samples of an 190 ka basanitic lava from the Eifel
volcanic field, Germany. The groundmass samples have
(230Th/238U) of 1.050 ± 0.015, whereas both the magnetite
and pyroxene separates lie within error of the equiline (i.e.,
230
Th in equilibrium with 238U) at lower (238U/232Th), so
that a tie line between either phase and the groundmass has
a slope greater than the equiline and thus is not a valid
isochron (Figure 10c). This indicates that the magnetite and
pyroxene crystals are older than the groundmass and are not
in chemical equilibrium with it and thus represent xenocrysts rather than phenocrysts. Returning to the example of
the zoned Laacher See eruption [Bourdon et al., 1994], in
contrast to the crystal-rich mafic phonolitic units that give
internal 230Th-238U mineral isochrons identical to the 13 ka
eruption age, glass and mineral separates from the crystalpoor felsic phonolite units show more scatter, with a poorly
defined isochron age of 30 ka (Figure 10b). This apparent
older age is interpreted as being due to inheritance of
crystals from an earlier episode of differentiation within
the magma chamber.
[66] A good example of nonconcordant 230Th-238U and
226
Ra-230Th age information and how this can give fresh
insights into the complex history of crystals is provided by
studies of historic lavas and cogenetic cumulate xenoliths
from Soufriere volcano (St. Vincent, Lesser Antilles
[Heath et al., 1998; Turner et al., 2003c]). Data on mineral
separates from a 1979 A.D. lava and from a cumulate
xenolith lie on a common array on the 238U-230Th equiline
diagram (Figure 12a) that corresponds to an age of 47 ±
10 ka, and yet all phases also preserve 226Ra-230Th
disequilibria. However, using the approach of Cooper et
al. [2001], 226Ra/Ba evolution curves for all minerals from
the lava and cumulate plot below that for the host lava
(Figure 12b), which demonstrates that none of the minerals
grew in equilibrium with the present groundmass. The
crystal size distribution of plagioclase grains for the host
lava [Turner et al., 2003c] indicates that some process has
caused a relative increase in the number of large crystals
and suggests a mixture between two crystal populations
that have grown at different rates in different environments. These data together suggest that the minerals are
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Figure 12. (a) The 230Th-238U equiline diagram for
separated phases from a lava (STV 354) and a cumulate
xenolith (WI 1A 18) from the 1979 A.D. eruption of
Soufriere volcano (St. Vincent, Lesser Antilles) that define a
common trend corresponding to an age of 47 (+12, 9) ka
[Heath et al., 1998; Turner et al., 2003c]. (b) The 226Ra/Ba
versus time evolution diagram for STV 354 whole rock and
calculated liquids in equilibrium with mineral separates
from STV 354 (solid curves) and WI 1A 18 (shaded curves)
[Turner et al., 2003c]. Abbreviations are cpx, clinopyroxene; pl, plagioclase; gm, groundmass; and ol, olivine.
Figure 12 is modified from Turner et al. [2003c], with
permission from Elsevier.
zoned in terms of composition and age, and the 226Ra
excesses are due to young overgrowths on old cumulate
crystals that were entrained by recent injection of a new
magma batch. The 230Th-238U apparent age indicates a
significant residence time for crystals in the cumulates
beneath the Soufriere volcano.
[ 67 ] Similar examples of discrepancies between
230
Th- 238 U and 226 Ra- 230 Th mineral ages have been
explained either by young growth of crystal rims on older
crystal cores (e.g., Tonga arc lavas [Turner et al., 2003c])
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or by mixtures of young and old crystals (e.g., Castle
Creek lavas at Mount St. Helens [Cooper and Reid,
2003]). Reagan et al. [1992] presented data showing
nonconcordant ages for the 226Ra-230Th and 228Ra-232Th
systems for anorthoclase crystal growth in phonolitic
magmas erupted at Mount Erebus (Antarctica) that can
also be explained by younger overgrowths on crystals.
The 226Ra-230Th-Ba data indicated an average age of
2400 years for anorthoclase growth, which is inferred
to have occurred within a shallow reservoir system as melt
inclusions within the crystals all have very low volatile
contents. Disequilibrium between 228Ra and 232Th in the
anorthoclase separates was attributed to growth of thin
rims on the crystals within 30 years of eruption, which
was probably caused by cooling associated with intrusion
into the lava lake present at the summit.
4.3.4. Effects of Crystallization of Th- and/or U-Rich
Accessory Phases on 238U-230Th Disequilibria
[68] The crystallization of accessory phases with very
high U and/or Th contents (e.g., zircon, apatite, sphene,
chevkinite, allanite, and monazite) and low modal abundance, particularly in felsic magmas, provides an additional
means of gaining fresh insights into the evolution of
magmatic systems. Most major phenocryst phases often
contain inclusions of these accessory phases that will
dominate the U and Th budget of the analyzed mineral
separate. Their presence can sometimes lead to analytical
problems as many accessory phases are difficult to dissolve
completely with conventional acid digestion techniques
[Condomines et al., 2003]. A mineral isochron determined
on the major phases may, in fact, be a mixing line controlled
by varying proportions of one or more accessory phase(s),
so that the isochron age is only giving the crystallization age
of the accessory phase(s). For example, the 238U-230Th
mineral isochron of 25 ± 10 ka for a comendite lava from
Olkaria, Kenya, is apparently governed by inclusions of
chevkinite, a phase with very high Th contents (1.5 wt %
ThO2) and very low (238U/232Th) (0.06) [Heumann and
Davis, 2002].
[69] It is more difficult to establish how the timing of
crystallization of such accessory phases fits into the crystallization history of the major fractionating phases and
hence with the magma differentiation process. Growth of
these phases can occur at different times because they are
not controlled by phase equilibria, as is the case for the
major phenocryst phases, but by saturation in the melt of a
specific essential structural component (e.g., Zr, zircon; P,
apatite; and Th, monazite), which depends on melt composition and temperature. Accessory phases can also crystallize as a result of local saturation adjacent to a growing
crystal of one of the major phenocryst phases [e.g., Bacon,
1989], and thus its growth may be contemporaneous with
that of the major phenocryst phase rather than reflecting the
timing of saturation in the overall magma reservoir.
[70] Accessory phases can also show evidence for
multiple growth stages, as demonstrated by Charlier and
Zellmer [2000], who analyzed three different size fractions
of zircons from the 26.5 ka Oruanui eruption in the Taupo
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Figure 13. The 230Th-238U equiline diagram showing
different zircon-whole rock model ages for different zirconsize fractions from the 26.5 ka Oruanui rhyolite eruption
from the Taupo volcanic zone, New Zealand. Ages are
calculated as two-point model ages relative to the whole
rock analysis. Data are age-corrected back to the time of
eruption, and so all model ages indicate time prior to the
eruption. The 230Th-238U data and the size distribution
of zircons can be used to constrain a mixing model
between old cores and younger rims. Figure 13 is modified
from Charlier and Zellmer [2000], with permission from
Elsevier.
volcanic zone (New Zealand). They found that the zirconwhole rock 238U-230Th data gave preeruption ages that
were different in each size fraction, with the older ages
given by the larger zircons (<63 mm, 5.5 ± 0.8 ka; 63–
125 mm, 9.7 ± 1.7 ka; and 125– 250 mm, 12.3 ± 0.8 ka)
(see Figure 13). These data can be explained by a
continuous zircon growth model over a period of 90 kyr,
but cathodoluminescence images showed that the zircon
crystals typically have euhedral cores and rims, and the
data can be modeled instead by mixing an older population of zircons (27 kyr at the time of eruption) with a
young zircon rim overgrowth that crystallized shortly
before eruption.
4.3.5. Evolution of Major Rhyolitic Systems From
238
U-230Th Disequilibria in Accessory Phases
[71] A key issue in understanding the volcanic hazards
associated with large rhyolitic magma systems is the length
of time that magma resides at shallow levels in the buildup
to major eruptions (>100 km3 magma). Several studies have
demonstrated the potential of U series disequilibria in
accessory phases to provide important time constraints in
major rhyolitic systems: Long Valley caldera, California,
United States [Reid et al., 1997; Heumann et al., 2002];
Toba caldera, Indonesia [Vazquez and Reid, 2004]; Yellowstone caldera, Wyoming, United States [Vazquez and Reid,
2002]; and Taupo volcanic zone, New Zealand [Charlier
and Zellmer, 2000; Charlier et al., 2003].
[72] Postcaldera rhyolites at Long Valley have eruption
ages of 110 ka based on 40Ar/39Ar dating of sanidine
crystals. However, Heumann et al. [2002] showed that
glasses from these units define a Rb-Sr isochron age of
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257 ± 39 ka (Figure 14a) that is interpreted as reflecting a
feldspar fractionation event 150 kyr before eruption.
(Note that extensive feldspar fractionation can produce very
large Rb/Sr ratios in high-silica magmas that are extreme
enough to produce resolvable variations in 87Sr/86Sr over a
period of a thousand years or so from the decay of 87Rb,
despite its long half-life of 48.8 109 years.) The
238
U-230Th data on mineral separates from one of the
postcaldera units (Deer Mountain) define a linear trend
corresponding to an age of 236 ± 1 ka (Figure 14b). As
all the major phases contain inclusions of zircon (high
U/Th) and allanite (low U/Th), this array represents a
mixing line between zircon and allanite populations. It
only has age significance if zircon and allanite crystallized
contemporaneously, and the observation that the glass
analysis plots below the array indicates that this cannot
be true. Instead, the data are consistent with two separate
and short (1 kyr) episodes of accessory phase crystallization, with zircon growth at 250 ± 3 ka, followed by
allanite growth at 187 ± 9 ka.
[73] Reid et al. [1997] developed an ion probe technique
to measure in situ 230Th-238U model ages of zircons in
rhyolites from Long Valley. Although individual model
ages are not very precise, this method allows many
crystals to be dated and thus provides an age spectrum
of a crystal population. For the Deer Mountain rhyolite,
except for a few young zircon ages close to that of the
eruption, most zircon ages were within error of a weighted
mean age of 226 ± 18 ka (1s), consistent with the thermal
ionization mass spectrometry data of Heumann et al.
[2002] (Figure 14c). A sample from a much younger
rhyolite (Inyo dome, 0.6 ka) from the same region had a
similar zircon age population with a weighted mean age of
229 ± 22 ka (1s). This is consistent with other evidence
that suggests that this eruption tapped the same magma
body as the Deer Mountain rhyolite 100 kyr earlier. Zircon
saturation temperatures indicate that the magmas cooled to
below 815C more than 200 kyr ago, but there is some
uncertainty over whether the magmas solidified then and
were subsequently remelted or whether they have been
crystal mushes since these zircons crystallized.
4.4. Overview of Differentiation and Crystallization
Timescales From U Series Disequilibria
4.4.1. Links Between U Series Residence Times
and Magma Composition and Eruption Rate?
[74] Hawkesworth et al. [2000] investigated how the
ages of crystal separates in recent volcanic rocks at the
time of eruption varied with an index of differentiation
(Figure 15a). Molar Si + Al was chosen as the index of
differentiation because these are framework-forming elements in the melt and there is a marked increase in the
viscosity of liquids at Si + Al = 66 [e.g., McBirney,
1984]. It was striking that most of the old preeruption
mineral ages are from rocks with relatively high Si + Al
values. Thus the simplest observation is the intuitive one
that the likelihood of erupted magmas containing old
crystals is much greater in the more evolved and hence
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higher-viscosity liquids. Moreover, simple cooling and
crystal-settling models imply that primitive basalts affected
solely by cooling in a chamber will contain a small
proportion of near-liquidus crystals, whereas more evolved
magmas will have an increasing proportion of older
crystals inherited from earlier stages of the magma’s
history.
[75] While the significance of old preeruption ages may
be different in different centers, it is striking that there is a
broad negative array between those ages and the average
eruption rates (Figure 15b). A marked exception is the
Auckland volcanic field where very small amounts of
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magma have been rapidly erupted from mantle depths,
which is where any differentiation took place [Rout et al.,
1993]. Nonetheless, in general, magmas with high eruption
rates tend to spend less time in the crust and vice versa, and
magmas that spend longer in the crust also tend to be more
evolved. Mid-ocean ridge basalts may take weeks to a few
thousand years to traverse the crust [Sigmarsson, 1996], and
it appears that volcanoes that erupt semicontinuously have
relatively small volumes of evolved magma types. In
contrast, evolved magma types tend to be more common
at volcanoes where magmas spend longer in the crust, the
eruptions are more episodic, and the volcanic center as been
active for 105 years. In western North America, for
example, there is typically 200 – 300 kyr of magmatic
activity at any center before caldera-related rhyolites are
erupted [Lipman, 1984; Reagan et al., 2003]. Moreover,
models of the thermal evolution of magmatic provinces
suggests that significant amounts of crustal melting will also
take place after 105 – 106 years at reasonable emplacement
rates of mafic magma [Annen and Sparks, 2002]. Another
inference of the broad array in Figure 15b is that it
constrains the rate of eruption, and hence presumably the
melt generation rate, of mafic magmas that traverse the
crust very rapidly, which may, in turn, inform models of
magma transport through the crust. Steady state models for
magmatic systems predict a negative relation between
magma residence times in a magma chamber (or in the
crust) and either the influx or outflux of magma, linked by
the volume of magma in the chamber (or in the crust) [e.g.,
Pyle, 1992]. Some of the time estimates in Figure 15b are
made on the basis of steady state modeling from Pyle
[1992], but for the others, there need be no simple link
between the estimated magma timescales and magma residence times. Nonetheless, it is interesting that steady state
modeling of the negative array in Figure 15b would suggest
Figure 14. Long Valley magmatic system, California,
United States. (a) Rb-Sr isochron diagram for postcaldera
glasses (eruption age 110 ka, from 40Ar/39Ar sanidine
ages) of an age 150 kyr older than the eruption age
(modified from Heumann et al. [2002], with permission
from Elsevier). (b) The 230Th-238U isochron diagram for
separated phases from the 110 ka postcaldera Deer
Mountain rhyolite [Heumann et al., 2002]. Amphibolebiotite-sanidine-zircon define a linear array corresponding
to an age of 236 ± 1 ka, but the glass and whole rock
samples plot below this array, close to the equiline. The
inset diagram shows a reduced scale to include the zircon
point with its high U/Th ratio. Figure 14b is modified
from Heumann et al. [2002], with permission from Elsevier.
(c) (238U/232Th) versus model 230Th-238U ages obtained by
in situ ion probe measurements of zircons extracted from a
sample of the Deer Mountain rhyolite [Reid et al., 1997].
Solid circles (with 1s error bars) show zircon data used to
calculate a weighted mean age of 226±18 ka (shown by
shaded rectangle). Three zircons (shaded circles) have
model ages similar to the eruption age. Figure 14c is
modified from Reid et al. [1997], with permission from
Elsevier.
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magma volumes in the crust of 10 –400 km3, consistent with
other estimates.
4.4.2. Links Between U Series Disequilibria Ages
and Earlier Volcanic Episodes?
[76] An important question is whether the U series
disequilibria data can aid our understanding of the behavior
of volcanoes. Several studies have highlighted intriguing
coincidences between time information obtained from U
series disequilibria data (e.g., preeruptive crystal ages and
whole rock magma differentiation ages) and specific episodes in the history of a volcanic system known from
surface stratigraphic studies (e.g., earlier eruptive episodes
and times when major changes in eruptive behavior
occurred). These include the following: (1) The Castle
Creek eruptive period 2000 years ago coincided with a
major shift in the eruptive behavior of the Mount St. Helens
volcano, from dominantly explosive dacitic volcanism to
dominantly effusive andesitic and basaltic lavas, and
Cooper and Reid [2003] were intrigued that many of the
plagioclase crystals entrained by younger lavas were
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apparently formed during this interval, based on U series
disequilibria data. (2) Hawkesworth et al. [2000] noted
that the 230 kyr differentiation timescale from basanite
to phonolite magma at Tenerife, inferred from variations in
(230Th/238U) in whole rock samples, is broadly similar to
the periodicity of the different eruptive cycles there over
the last 1.6 Myr. (3) Turner et al. [2003b] compared
226
Ra/230Th data from two adjacent and compositionally
similar volcanoes in Indonesia in the Sunda arc. Sangeang
Api lavas and cumulates indicated residence times of
2000 years in a relatively small 10 km3 magma
chamber, whereas a lava sample from the cataclysmic
1815 A.D. eruption (100 km3) from Tambora indicated a
longer crustal residence time of 5000 years that is identical
to the time elapsed since the preceding major eruption at
Tambora. Such speculative links to specific events in a
volcano’s history clearly merit further investigation at other
volcanoes with well-constrained stratigraphic records.
4.4.3. Summary
[77] It is clear that U series disequilibria data on melt and
crystal phases of igneous rocks can provide valuable temporal information on different stages in the magmatic
history of a particular volcanic center. The groundmass of
different related samples can be used to estimate the
timescale for a parental magma to differentiate to more
evolved compositions, and where the effects of magma
mixing and crustal assimilation appear to have been minimal, estimates range from 1000 years [e.g., Vigier et al.,
1999] to 230,000 years [e.g., Hawkesworth et al., 2000].
In general, the ages indicate slower differentiation rates at
the higher temperatures in the deep crust and more rapid
differentiation at cooler shallow crustal levels. Crystals
found in igneous rocks can provide a complex record of
magmatic processes, preserved as both textural and compositional variations, and yet it is often not clear how these
variations are related to the evolutionary history of the host
Figure 15. (a) Preeruption ages of various recent magmatic suites as derived from U series disequilibria versus an
index of differentiation, in this case molar Si + Al of the
bulk rock [Hawkesworth et al., 2000]. Most old 238U-230Th
mineral isochrons have been obtained from rocks with
relatively high Si + Al contents and higher viscosities. By
contrast, studies of bulk rock variations of 238U-230Th in
suites of comagmatic but compositionally diverse samples
suggest that differentiation from mafic parental magmas to
more intermediate compositions can take 10– 200 kyr.
Similar studies, principally using whole rock 226Ra-230Th
data, suggest that the inferred differentiation times for more
evolved magmas are significantly shorter, perhaps tens to
hundreds of years. Overall, these observations are consistent
with slower rates of differentiation in more mafic magmas at
greater depths and much faster rates of differentiation in
more evolved magmas at shallower depths. (b) Variation of
eruption rates [from Crisp, 1984] with preeruption ages
estimated for a wide range of different magmatic suites (see
Hawkesworth et al. [2004] for details). Figure 15 is taken
from Hawkesworth et al. [2004], with permission from
Elsevier.
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magma. Petrographic observations often show clear evidence for the complex multistage growth history of crystals,
with overgrowths of young rims on older inherited cores
or xenocrysts perhaps related to an earlier stage in the
magmatic system and patterns of zoning and resorption
indicating fluctuating pressure-temperature-fO2 conditions
during crystal growth over a time period that is long
relative to the nuclide half-life. It is important to realize
that U series analysis of mineral separates will tend to
average such effects, and in addition, the U, Th, and Ra
budgets of bulk mineral separates might be dominated by
inclusions of trapped melt or accessory phases. The recent
study by Vazquez and Reid [2004], who measured trace
element and 230Th-238U variations in situ in zoned allanite
crystals from the voluminous 75 ka Toba Tuff, represents
an important advance because the trace element variations
can be linked to specific stages in the differentiation of the
magma at the time given by 230Th-238U data. In conclusion, U series disequilibria data have provided important
constraints on the timing of crystal growth and residence
time in magmas.
5. TIMESCALES AND PROCESSES OF MELT
GENERATION AND TRANSPORT
5.1. Background to Melt Generation Processes in
Upwelling Mantle
[78] Melt generation processes are best constrained under
mid-ocean ridges, where partial melting occurs in response
to decompression in the upper mantle as the plates move
away from the spreading center and new crust is generated
in their place. The application of trace element data greatly
increased the sensitivity of partial melting and fractional
crystallization models [e.g., Gast, 1968], and initially at
least, these models assumed bulk equilibrium between the
melt and the residual matrix until the magma was extracted
(i.e., batch melting). U and Th are highly incompatible
elements in the upper mantle, and so U/Th ratios in the melt
should be similar to those in their source [O’Nions and
McKenzie, 1993; Elliott, 1997] at the large degrees of
partial melting (10%) inferred for MORB [e.g., Klein
and Langmuir, 1987; McKenzie and Bickle, 1988]. Thus
the expectation was that 238U and 230Th should be close to
secular equilibrium in MORB magmas.
[ 7 9 ] However, the demonstration of significant
238
U-230Th isotope disequilibria in MORB [Condomines
et al., 1981; Newman et al., 1983] highlighted that such
models might be overly simplistic [e.g., McKenzie, 1985].
Element fractionation is only efficient when the degree of
melting is similar to or less than the bulk distribution
coefficient between melt and mantle minerals, which for
U and Th is <0.01. Thus, to obtain 238U-230Th disequilibria,
the degrees of melting must be small (<1%), and so
dynamic melting models were developed in which small
degree melts migrated through the melt zone and were
then mixed together with larger degree melts before melt
extraction [e.g., Williams and Gill, 1989]. In such models
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the degree of U series isotope disequilibria depends both
on the porosity in the melt generation zone, and hence the
fraction of melt in equilibrium with the solid, and on the
melting rate relative to the half-life of the daughter
product. Thus the duration of the melting process is
important, and as melting rate depends on potential temperature and the upwelling rate, there is a direct link to
mantle dynamics. In the relatively simple tectonic setting
of mid-ocean ridges it is possible to estimate the solid
mantle upwelling rate directly from the plate separation or
spreading rate.
[80] In the dynamic melting model [McKenzie, 1985], as
melt is generated it remains in equilibrium with the
residual mantle until a threshold porosity is reached, at
which point any melt in excess of this porosity value is
instantaneously extracted to the surface in chemically
isolated channels without further interaction with residual
mantle (Figure 16a). This dynamic melting model may
now be usefully regarded as one end-member of a number
of open system ‘‘ingrowth models’’ [Williams and Gill,
1989; Spiegelman and Elliott, 1993] in which excess
daughter products over that in the source are produced
by ingrowth from the parent because of their different
residence times within the melting column. For example,
peridotite residual after extraction of a small degree melt
will have an elevated U/Th ratio and plot to the right of
the equiline, and subsequent radioactive decay will result
in higher 230Th/232Th ratios that may then be mobilized by
further melting (Figure 16c). For equilibrium dynamic
melting, significant radioactive disequilibrium can therefore be generated by ingrowth provided that the melting
rate is slow relative to the decay rate of the short-lived
daughter isotope (e.g., slow upwelling mantle or a long
melting column so that the duration of melting is significant). For these ingrowth models, (226Ra/230Th) is controlled mostly by the porosity of the melt region (which
controls the velocity of the melt relative to the residual
mantle), whereas (230Th/238U) and (231Pa/235U) are controlled more by the melting rate (which is related to the
residual mantle upwelling velocity). In practice, the inferred porosities are small relative to those estimated from
modeling of trace element abundances in abyssal peridotites and melt inclusions [e.g., Johnson et al., 1990; Slater
et al., 2001]. One explanation may be that the melt is in
equilibrium with the crystal surfaces, and as crystal-liquid
interactions are sensitive to the diffusion of elements
within the crystals [Qin, 1992, 1993], this will change
the effective partition coefficients and allow larger inferred
porosities.
[81] The other end-member ingrowth model involves
melt generation and continuous chemical equilibrium
between melt and residual mantle during melt percolation
and transport (e.g., equilibrium or chromatographic or
reactive porous flow model (Figure 16b) [Spiegelman and
Elliott, 1993]). This equilibrium porous flow affects the
relative transport velocities and hence the overall residence
times of the parent and daughter nuclides in the melt
column. In single porosity models the melt velocity
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depends on the solid upwelling velocity, the degree of
melting, and the porosity at the same height. The porosities
are still small (0.2% for grain sizes of 3 mm), but these
can produce significant U series disequilibrium provided
that the parent spends longer in the system than the
daughter [Spiegelman and Kelemen, 2003]. Spiegelman
and Elliott [1993] also discussed other possible mechanisms that might produce disequilibria as a result of
differential transport velocities of parent and daughter
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nuclides, such as different degrees of surface adsorption.
Hiraga et al. [2004] presented a thermodynamic model
for equilibrium grain boundary segregation that predicts
that grain boundaries should preferentially host incompatible elements such as U, Th, and Ra, thus providing
another potential mechanism for their selective enrichment
in melts that, as yet, has not been considered in melting
models.
[82] For a given set of model parameters, equilibrium
porous flow models will generate larger disequilibria than
dynamic melting models because the difference in residence
time in the melting column between parent and daughter is
maximized [Spiegelman and Elliott, 1993]. The major
difference between dynamic melting models and equilibrium porous flow models is where the U series disequilibria are set up within the melting column. In dynamic
melting the U series disequilibria are generated during the
onset of melting near the bottom of the melting column,
and so preservation of 226Ra excesses in magmas at the
surface requires very rapid rates of melt transport. In
equilibrium porous flow models the U series disequilibria
are created throughout the melting column, and this
Figure 16. Cartoons to illustrate the main features of the
two end-member ‘‘ingrowth’’ melting models [from
Spiegelman and Elliott [1993]. (a) Dynamic melting
[McKenzie, 1985], where each instantaneous melt increment
is extracted, transported instantaneously in chemical isolation, and then mixed with the other instantaneous melts
produced at other depths in the melting column prior to
eruption. Incompatible elements are effectively removed
from the melting region during the initial stages of melting
(where the degree of melting is similar to the bulk partition
coefficients of the nuclides of interest), and so nearly all of
the daughter ingrowth takes place at the base of the melting
column and requires rapid extraction to be preserved in an
eruption at the surface. (b) Equilibrium porous flow
[Spiegelman and Elliott, 1993], where the melt interacts
with residual matrix mantle during transport to the surface. If
the daughter nuclide is more incompatible than the parent
nuclide, it will have a faster effective velocity through the
melt column, and so ingrowth of additional daughter
nuclides will take place because the parent nuclide spends
more time in the melting column than the daughter nuclide.
(c) The principle of nuclide ‘‘ingrowth’’ is illustrated by the
U-Th equiline diagram [Spiegelman and Elliott, 1993],
which shows how the activities of 238U and 230Th change in
both the melt and the residual solid mantle as the degree of
melting increases from 0% (base of melting column) to 25%
(top of melting column). Elemental fractionation of U
and Th is only significant at low degrees of melting. If
DU > DTh, as in this example, then the solid becomes
enriched in 238U, and as this decays, the solid and any
subsequently formed melt will have higher 230Th/232Th, thus
producing ‘‘ingrowth’’ of 230Th. The key parameters that
will influence the disequilibria values for (230Th/238U),
(226Ra/230Th), and (231Pa/235U) are source mineralogy,
partition coefficients within the melting column, matrix
porosity, mantle upwelling velocity, melt transport rate,
melt-mantle interaction during transport, and presence of
volatiles.
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Figure 17. (a) (230Th/238U) versus axial ridge depth,
showing a broad negative correlation of decreasing 230Th
excesses with increasing water depth (updated from
Bourdon et al. [1996b] by Lundstrom [2003]). (b) Slope
of segment-scale linear data trends on the 230Th-238U
equiline diagram versus half-spreading rates of the ridge
segment [Lundstrom et al., 1998a; Lundstrom, 2003].
Abbreviations are as follows: AAD, Australian-Antarctic
discordance; EPR, East Pacific Rise; FAZAR, 33N –40N
MAR; JdF, Juan de Fuca; KR, Kolbeinsey Ridge; MAR,
Mid-Atlantic Ridge; and RR, Reykjanes Ridge. Figure 17 is
taken from Lundstrom [2004], with permission from
Mineralogical Society of America.
lessens any requirement for rapid melt ascent to explain
large 226Ra excesses. The latest models for melt generation
beneath ridges have focused on hybrid models that combine aspects of these two end-member models. In these
‘‘two-porosity’’ models, melts migrate in channels that are
fed by porous flow on the grain scale in the country rock
[Spiegelman and Elliott, 1993; Iwamori, 1994; Kelemen et
al., 1997; Lundstrom, 2000; Jull et al., 2002]. Such
models increase the transport rate of 230Th from depth
and allow isotope disequilibria to be generated at different
depths.
5.2. Mid-Ocean Ridge Basalts
5.2.1. Observations
[83] Obtaining comprehensive U series disequilibria data
on suites of MORB glasses is not straightforward, not least
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because the eruption ages of most samples are unknown.
This means that the 226Ra-230Th disequilibria is used
primarily to confirm young (<8 kyr) ages for lava flows
rather than in constraining melting processes. Furthermore,
most sample collection is done by dredging, and sample
suites tend to be regional in coverage rather than focusing
on local segment-scale variations, although use of submersibles can improve sample density and help to target
fresh-looking flows [e.g., Sims et al., 2002a]. Samples also
tend to have lower trace element abundances compared to
lavas from other tectonic settings, and there is greater
potential for seawater alteration and contamination by
Fe-Mn crusts (see section 2.3), which all conspire to make
it analytically more challenging to obtain high-quality U
series disequilibria data.
[84] The 230Th-238U disequilibria data can be used as a
tool for dating mid-ocean ridge basalts where eruption
ages are otherwise difficult to determine [e.g., Goldstein et
al., 1992]. Young MORB lavas (essentially those with
226
Ra-230Th disequilibrium, i.e., <8000 years) from local
ridge segments, in general, form subhorizontal arrays on the
U-Th equiline diagram (Figure 18a) [e.g., Lundstrom et al.,
1998a]. Provided that the melt generation process in the
region has remained uniform over the last 100,000 years
or so, such arrays can be used to estimate the initial
230
Th/232Th of any MORB sample in the region from its
measured U/Th ratio. Off-axis samples generally have
lower than expected 230Th/232Th, and the simplest explanation is that this difference reflects posteruption radioactive decay of 230Th and can therefore be used to calculate
a model eruption age for each sample.
[85] There are five global-scale observations from U
series disequilibria measurements of MORB lavas that need
to be discussed and explained: (1) the almost ubiquitous
230
Th excesses (Figures 6 and 7); (2) the elevated
(231Pa/235U) compared with within-plate and subduction
zone magmas (Figure 7b) [e.g., Pickett and Murrell,
1997]; (3) the broad correlation of (230Th/238U) with
axial ridge depth (Figure 17a) [e.g., Bourdon et al.,
1996b]; (4) the apparent correlation with spreading rate
of the slope of data trends from different ridge segments
on the 238U-230Th equiline diagram (Figure 17b) [e.g.,
Lundstrom et al., 1998a]; and (5) the negative correlation
between 226Ra excess and 230Th excess (Figure 18c)
[Kelemen et al., 1997; Sims et al., 2002a]. Recent reviews
by Lundstrom [2003] and Elliott and Spiegelman [2003]
provide extensive details of the observed variations of U
series disequilibria in MORB samples and the potential
constraints such data can provide on the nature of the
melt generation and transport processes beneath ridges.
5.2.2. Constraints on Source Mineralogy
and Depth of Melting
[86] The 230Th excesses found in most MORB magmas
imply that Th is more incompatible than U during mantle
melting, i.e., bulk DTh/DU < 1, so that U is preferentially
retained in the residual solid relative to Th. Partition
coefficient data for U and Th in mantle minerals (summarized by Blundy and Wood [2003]) indicate that the 230Th
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Figure 18. The best studied mid-ocean-ridge segment is the East Pacific Rise at 8 – 10N (bounded by
the Siqueiros Transform to the south and the Clipperton Transform to the north), which has an abundance
of samples with known young ages (<8 ka) [Goldstein et al., 1993; Volpe and Goldstein, 1993;
Lundstrom et al., 1999; Sims et al., 2002a]. (a) The 230Th-238U equiline diagram, showing a broad linear
correlation. (b) Th content versus (238U/232Th), showing that ‘‘depleted’’ samples with low Th have high
(238U/232Th) and low (230Th/238U). (c) (226Ra/230Th) versus (230Th/238U), showing an inverse correlation.
(d) (231Pa/230Th) versus (230Th/238U), showing little correlation.
excesses require small degree melting in the presence of
garnet [Beattie, 1993; LaTourrette et al., 1993] or highpressure aluminous clinopyroxene [Landwehr et al.,
2001], as these phases both have DTh/DU < 1 and DU
of 102 – 103. U and Th have extremely low mineralmelt partition coefficients in olivine and orthopyroxene
with DU of 104 – 105, at least an order of magnitude
smaller than for garnet or clinopyroxene, giving them
little leverage over U-Th fractionation, and so neither
olivine nor orthopyroxene have major roles to play in
governing 230Th-238U disequilibria. Clinopyroxene is a
special case as the relative partitioning of U and Th is
very sensitive to composition, which varies as a function
of temperature and pressure. At shallow mantle pressures,
DTh/DU > 1, but with increasing depth the clinopyroxenes
become more aluminous with DTh/DU < 1 at depths
greater than 50 km [Landwehr et al., 2001]. Garnet is
stable at depths >75 –90 km [Robinson et al., 1998; Sims
et al., 2002a]. Thus 230Th excesses may be produced
during melting at pressures greater than 1.5 GPa (i.e.,
>50 km), irrespective of the presence of garnet. However,
garnet is much more efficient at generating 230Th excesses
because DTh/DU is much less in garnet than in clinopyroxene [Landwehr et al., 2001]. Although partition coefficients for Ra and Pa have not been measured directly for
most mantle minerals, it seems reasonable to assume that
they are very small, such that DU > DPa and DTh > DRa
for both garnet peridotite and spinel peridotite [e.g.,
Blundy and Wood, 2003]. This is consistent with the
observation that MORB lavas have (231Pa/235U) > 1 and
(226Ra/230Th) > 1.
[87] Although rare, a few examples of MORB lavas
with (230Th/238U) < 1 have been found [e.g., Bourdon et
al., 1996b; Sims et al., 2002b; Tepley et al., 2004]. Tepley
et al. [2004] showed that the measured disequilibria in
depleted samples from the Garrett Transform with 238U
excesses (accompanied by significant 231Pa and 226Ra
excesses) and low U and Th abundances could be
modeled by shallow melting of depleted spinel lherzolite.
Furthermore, Lundstrom [2000] noted that MORB samples with 238U excesses generally have major element
compositions similar to experimental melts of spinel
lherzolite at 1 GPa.
[88] As shown on Figure 7b, MORB magmas have
greater 231Pa excesses than both within-plate magmas
and subduction-related magmas. Most within-plate magmas are generated beneath a relatively thick lithosphere
that restricts the height of the melt column so that most of
the melting takes place in the garnet peridotite stability
zone. MORB melting, on the other hand, can continue to
shallower levels within the spinel peridotite stability zone,
and this longer melting column allows greater ingrowth of
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Pa [Lundstrom, 2003]. It is notable that the within-plate
magmas that plot closest to the MORB field are from the
Asal rift and Iceland [Pickett and Murrell, 1997; Vigier et
al., 1999], both in extensional tectonic settings more
similar to a ridge environment. The 231Pa and 230Th
excesses in MORB can be explained by shallower melting
of mantle with a higher clinopyroxene:garnet ratio than
most ocean island basalts [Bourdon et al., 1998; Bourdon
and Sims, 2003].
5.2.3. Influence of Variations in Mantle Temperature,
Spreading Rate, and Source Heterogeneity
[89] Global correlations between major element compositions of MORB and axial depth of the ridge have been
explained by along-axis mantle temperature variations
[Klein and Langmuir, 1987], and this model can also
account for the broad negative correlation between
(230Th/238U) and axial depth [Bourdon et al., 1996b].
Higher mantle temperatures mean that melting initiates at
a greater depth, resulting in a longer melt column and more
melt generation, which, in turn, forms a thicker crust and a
shallow depth for the ridge. Bourdon et al. [1996b] quantitatively explained the increase in 230Th excess with
increasing depth of melt initiation within the garnet stability
field, using an equilibrium porous flow model. Nonetheless, there can be a wide range in (230Th/238U) within
particular area and depth ranges (e.g., Juan de Fuca
samples have a range in (230Th/238U) from 1.1 to 1.4 at
constant depth of 2200 m), and this is usually interpreted
as primarily the result of local source heterogeneities
[Bourdon et al., 1996b; Lundstrom et al., 1998a].
[90] MORB samples from particular regions or segments
nearly always show a marked correlation between
(230Th/238U) and U/Th. A good example is provided by
data from the well-studied 8– 10N region of the East
Pacific Rise (EPR) (Figure 18a). In detail, the slope of the
trends vary (Figure 17b), and Lundstrom et al. [1998a]
argued that they are correlated with the half-spreading rate
of the particular ridge. In contrast, Elliott and Spiegelman
[2003] argued that these covariations are not significant.
Additional studies where there are better constraints on the
slopes of local data sets are required to resolve this issue,
particularly for those ridge segments for which only older,
a-counting data are presently available. Furthermore, the
primary mechanism behind the generation of these linear
trends is still the subject of much debate. Melt generation
beneath ridges is a polybaric process with melt produced
from a range of depths within a variably depleted mantle
and then mixed prior to eruption. The linear trends are
consistent with mixing between two melts derived from
different depths in the melting column, but the ultimate
origin of the end-member melts is not clear. Lundstrom et
al. [1998a] interpreted these local trends as being due to
binary mixing of melts derived from ‘‘enriched’’ and
‘‘depleted’’ sources within a heterogeneous mantle upwelling at the same rate but melted at different depths, whereas
Sims et al. [2002a] and Elliott and Spiegelman [2003]
argued that complexities of the melt generation, segregation,
transport, and mixing process are capable of generating
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melts from a homogeneous source with a range in U/Th
values.
[91] Several ridge segments display significant local
variations in U/Th coupled with limited variations in the
compositions of long-lived radiogenic isotopes such as
87
Sr/86Sr, 143Nd/144Nd, 176Hf/177Hf, and 208Pb/204Pb (e.g.,
9– 10N EPR [Sims et al., 2002a] and Kolbeinsey Ridge
[Sims et al., 2002b]), and Sims et al. [2002a, 2002b]
therefore argued that the mantle source must essentially
be homogeneous in both areas and that polybaric melting
and progressive source depletion rather than source heterogeneity produced the observed variations in U/Th. However, the significant diversity in trace element and
radiogenic isotope compositions of off-axis lavas [Niu and
Batiza, 1997; Niu et al., 2002] suggests that the mantle in
the EPR area is heterogeneous on a length scale smaller than
ridge segments. Lundstrom [2003] has also argued that
mixing curves between the depleted and enriched endmembers are strongly hyperbolic such that limited variations in, for example, 87Sr/86Sr that are linked to large
variations in U/Th can still be consistent with source
heterogeneities. Clearly, additional comprehensive data
sets on suites of local MORB samples that combine
long-lived radiogenic isotope measurements and trace
element analyses with U series disequilibria data are
required to ascertain the relative roles of source heterogeneity and melting processes in generating local variations in
U/Th. Application of high-precision Pb isotope measurements should be particularly beneficial as they offer greater
resolution of subtle compositional effects due to mixing of
melts from heterogeneous sources [Galer et al., 1999;
Thirlwall et al., 2004].
5.2.4. Constraints From (226Ra/230Th) Disequilibria
[92] As mentioned in section 5.1, determining the depth
of generation of the observed 226Ra-230Th disequilibria in
MORB can place important constraints on the nature of the
melt migration process (rapid melt transport through conduits or slow melt percolation during equilibrium porous
flow) because in dynamic melting models all disequilibria
are generated at the base of the melting column, whereas in
equilibrium porous flow models different disequilibria can
be generated at different levels within the melt column.
Samples with known eruption ages from the East Pacific
Rise (8– 10N) and the Juan de Fuca Ridge show an
inverse correlation between 226Ra excess and 230Th excess
(Figure 18c). This observation is inconsistent with a purely
dynamic melting model, which would predict a positive
correlation between 226Ra excess and 230Th excess [Sims et
al., 2002a]. On the other hand, the observations that the
major and trace element compositions of MORB melts are
not in equilibrium with the shallow upper mantle [Kelemen
et al., 1997] and that abyssal peridotites represent residues
from near-fractional melting processes [Johnson et al.,
1990] together indicate that purely equilibrium porous flow
models cannot be the only process involved in the generation and transport of MORB melts.
[93] Models have been developed that combine different
porosity regimes in the mantle melting column [Kelemen et
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al., 1997; Lundstrom, 2000; Lundstrom et al., 2000; Jull et
al., 2002; Sims et al., 2002a]. In these ‘‘two-porosity’’
models the anticorrelation between 226Ra excess and
230
Th excess is produced by mixing of slow moving melt
transported through reactive, low-porosity mantle with
relatively fast moving melt transported within unreactive,
high-porosity channels, again emphasizing the role of melt
mixing in the generation of MORB magmas. One problem
with these models has been the difficulty of deriving sets of
parameters that can simultaneously produce the observed U
series disequilibria and still match the extreme elemental
depletions observed in abyssal peridotites. An important
caveat to this discussion is that these interpretations all
assume that the 226Ra-230Th disequilibrium results from the
melting process. Saal and Van Orman [2004] recently
suggested that 226Ra excesses might result from diffusive
interaction between melts and shallow cumulates. More data
are required from other mid-ocean ridge segments to test the
robustness of the (230Th/238U)-(226Ra/230Th) anticorrelation
observed at the Juan de Fuca Ridge and East Pacific Rise.
[94] Fluid dynamical models for reactive flow in a
deformable permeable matrix [Aharonov et al., 1995;
Spiegelman et al., 2001] have shown that a coalescing
network of high-porosity melt channels surrounded by
extremely low porosity regions will develop in the mantle
melt column, consistent with field observations of reactive
dunite channels in the mantle sections of ophiolites [e.g.,
Kelemen et al., 2000]. Modeling of the chemical consequences of such full reactive transport models indicates that
significant chemical diversity will be produced in melts of a
homogeneous source, consistent with the observed variability in melt inclusions trapped in MORB olivine crystals
[Spiegelman and Kelemen, 2003]. More importantly, preliminary investigations [Elliott and Spiegelman, 2003]
indicate that these models are capable of at least qualitatively reproducing some of the critical U series disequilibria
observations, namely, the linear arrays of (238U/232Th)
versus (230Th/ 232 Th) and the negative correlation of
(230Th/238U) with (226Ra/230Th), from melting of a homogeneous source, and they clearly merit further attention.
5.3. Within-Plate Magmatism
5.3.1. Background
[95] In this section we discuss both oceanic island basalts
and continental intraplate lavas. Within-plate magmas are
compositionally more heterogeneous than MORB, both on a
global scale and on a local scale. This is particularly
apparent both from variations in radiogenic isotopes
(Sr-Nd-Pb-Hf-Os) and elemental concentrations and indicates a wider diversity of potential mantle source materials
and degrees of melting [e.g., Zindler and Hart, 1986;
Hawkesworth et al., 1990; Hofmann, 1997]. Physical
details of the melt generation process are less certain than
at mid-ocean ridges, and it is not as easy to constrain
uniquely the degrees of melting or to ascertain the links
to geodynamical parameters such as mantle upwelling
velocity. The presence of the overlying lithosphere adds
additional complexity to the melt generation process by
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controlling the height of the melting column and also by
acting as a potential source of melts. Bourdon and Sims
[2003] provide a comprehensive review on the constraints
that U series disequilibria can provide on the nature of
melt generation processes in within-plate settings.
5.3.2. Relative Roles of Elemental Fractionation
and Nuclide Ingrowth on Disequilibria
[96] As emphasized by Elliott [1997], a key issue for
understanding U series disequilibrium in within-plate
basalts (especially OIB) is determining the relative effects
of net elemental fractionation versus nuclide ingrowth. The
high degrees of melting (10%) inferred for the generation
of MORB magmas and the observed large 230Th and 231Pa
excesses are consistent with an ingrowth model and slow
melting rates. However, many within-plate lavas have
compositional features (enriched incompatible trace element
patterns and silica undersaturation) that indicate much lower
degrees of melting [Green and Ringwood, 1967; Gast,
1968] where net elemental fractionation has a greater
potential to influence U series disequilibrium.
[97] If 230Th-238U disequilibrium is established by net
elemental fractionation via a batch melting process, then the
measured (230Th/232Th) in a magma should be equivalent to
the (238U/232Th) and hence a robust estimate of the U/Th of
the mantle source, provided melt production and extraction
is rapid. If the melting rate is slow, then ingrowth of 230Th
in the residual mantle becomes significant, causing the
(230Th/232Th) in the magma to be greater than in the source.
Thus, if ingrowth is significant in controlling 230Th-238U
disequilibrium in most within-plate basalts, then the U/Th
of the magma should be a better estimate of the source
U/Th. If variations in the U/Th ratios of mantle sources are
long-lived and related to elemental fractionations of other
elements such as Rb-Sr, then correlations between 87Sr/86Sr
and an estimate of source U/Th would be expected. Recent
compilations of OIB data by Condomines and Sigmarsson
[2000] and Bourdon and Sims [2003] show better correlations between 87 Sr/ 86 Sr and ( 230 Th/ 232 Th) than with
(238U/232Th), and this has been used to infer that for many
OIB lavas, (230Th/232Th) is a better estimate of source U/Th
and that net elemental fractionation of U and Th can be
important. However, the diverse range in radiogenic isotope
compositions shown by OIB magmas requires at least four
types of distinct mantle source components [e.g., Hofmann,
1997], and it is not clear that any robust correlation between
87
Sr/86Sr and either (230Th/232Th) or (238U/232Th) would be
expected.
[98] Sims et al. [1995, 1999] showed that the extent of
230
Th-238U and 231Pa-235U disequilibrium in Hawaiian lavas
(Figures 19a and 19d) was related to the degree of melting
inferred from major and trace element systematics, with
small 230Th and 231Pa excesses (2 – 6% and 10 – 15%,
respectively) in tholeiites (large degree melts) and larger
230
Th and 231Pa excesses (15 – 30% and 78%, respectively)
in alkali basalts (small degree melts). Thus the 230Th-238U
and 231Pa-235U disequilibrium and rare earth element
variations could be explained by a batch melting model
of a garnet-bearing source (to create the 230Th excesses).
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Figure 19. Variations of U series disequilibria with distance from inferred center of hot spot upwellings
at Hawaii and Iceland: (a) (230Th/238U), (b) (238U/232Th), (c) (226Ra/230Th), and (d) (231Pa/235Th). Data
sources for Hawaii are Cohen and O’Nions, [1993], Cohen et al. [1996], Sims et al. [1999], and
Pietruszka et al. [2001]; data sources for Iceland are Peate et al. [2001a] and Kokfelt et al. [2003].
However, the significant 226Ra excesses in the same lavas
(Figure 19c) cannot be explained by such a model and
require some ingrowth of 226Ra during melting [Elliott,
1997; Sims et al., 1999].
5.3.3. Source Heterogeneities
[99] One issue in understanding the compositional diversity of OIB lavas is whether this is related to lithological
variations within the mantle source materials. Different
lithologies would have different mineral compositions and
hence different bulk partition coefficients for trace elements that might have a significant influence on the extent
of U series disequilibria developed in suites of OIB lavas.
In particular, there has been much discussion over the
possible role for garnet pyroxenite or eclogite veins
(approximately 70%– 50% clinopyroxene and 30%– 50%
garnet) in the source of OIB lavas. These lithologies are
observed in mantle xenolith suites and in peridotite
massifs, hosted in a peridotitic matrix (approximately
50% olivine, 20% orthopyroxene, 20% clinopyroxene,
and 10% garnet or spinel), and they are thought to
represent remnants or melts of recycled oceanic crust
[e.g., Hirschmann and Stolper, 1996; Sigmarsson et al.,
1998a; Stracke et al., 1999].
[100] It is an important issue because the ubiquitous
230
Th excesses in OIB lavas are generally assumed to be
a garnet signature, indicative of deep melting within the
garnet peridotite stability field. However, garnet-bearing
pyroxenites or eclogites are stable at shallower pressures
than garnet peridotite, and if they are responsible for the
230
Th excesses, then this lessens the estimates of the depth
of initiation of melting that can be incorporated into
geodynamic models. Different authors have reached different conclusions about the relative importance of pyroxenites in controlling 230Th-238U disequilibria in OIB [e.g.,
Sigmarsson et al., 1998a; Stracke et al., 1999]. The
mineral compositions of garnet and clinopyroxene in
pyroxenitic versus peridotitic lithologies are quite different,
and the compositional dependence of mineral/melt partition
coefficients should thus lead to different partitioning
behavior for U and Th, as well as other trace elements
such as Lu and Hf [e.g., Stracke et al., 1999; Pertermann
et al., 2004]. Therefore, provided we have the relevant
partitioning data and comprehensive trace element and U
series disequilibria analyses from a suite of OIB lavas, this
should be a resolvable issue.
5.3.4. Constraints on Variations in Mantle Upwelling
Rates Near Hot Spots
[101] Iceland is a ridge-centered hot spot, and yet the rift
tholeiites have lower than expected 230Th excesses based on
the global MORB correlation between 230Th excess and
ridge depth (Figure 17a). These tholeiites are generated by
large degrees of melting broadly similar to MORB magmas,
and so nuclide ingrowth is likely to be important in
controlling 230Th-238U disequilibrium. The lower than
expected 230Th excesses suggest faster melting rates, and
this is most simply explained as active mantle upwelling
(i.e., upwelling faster than expected from passive plate
separation), consistent with models for a hot, buoyant
mantle plume beneath most of Iceland [Bourdon et al.,
1996b; Peate et al., 2001a; Kokfelt et al., 2003]. Sleep
[1990] and Ribe [1996] used geophysical observations to
determine the buoyancy flux beneath many oceanic hot
spots. There is a direct link between buoyancy flux and
mantle upwelling velocity such that the inferred global
range in buoyancy flux equates to a threefold to fourfold
variation in upwelling velocity, which can explain the range
in observed 230Th and 231Pa excesses in ocean island basalts
[e.g., Bourdon et al., 1998]. The broad negative correlation
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between average (230Th/238U) at an oceanic hot spot and
buoyancy flux was noted by Chabaux and Allègre [1994]
and can be explained by variations in upwelling rate, with
the low (230Th/238U) values typical of tholeiites from the
high-buoyancy Hawaiian hot spot consistent with limited
230
Th ingrowth during fast mantle upwelling.
[102] Within individual hot spot locations, there are often
significant variations in U series disequilibria that will have
been influenced by several factors, including compositional
heterogeneities within the source, different extents of melting, and variable upwelling rates. In several cases the U
series disequilibria vary systematically with distance from
the inferred hot spot center (e.g., Hawaii [Sims et al., 1999],
Iceland [Kokfelt et al., 2003], and the Canaries [Lundstrom
et al., 2003]) (Figure 19), which is consistent with models
for radial variations in mantle upwelling rate above a mantle
plume: The lowest (230Th/238U) and (231Pa/235U) values are
found in lavas from the fast upwelling central axis, with
higher values found in the slower upwelling periphery
allowing more time for ingrowth. At least in the case of
Iceland the increase in 230Th excesses with increasing
distance from the inferred plume axis is found in samples
with similar U/Th ratios (Figure 19b), thus minimizing the
potential effects of source heterogeneities that might otherwise influence the magnitude of 230Th-238U disequilibria
through variations in melting rates (i.e., more/less productive mantle sources) or mineralogy. Differences in 230Th
excesses for a given U/Th ratio between samples from the
periphery of Iceland and the adjacent Reykjanes Ridge can
be interpreted as indicating that the whole of Iceland is
influenced by active upwelling [Peate et al., 2001a].
Although the absolute values for the solid upwelling
velocity beneath the different Hawaiian islands inferred
from the U series disequilibrium data are highly sensitive
to the choice of melting model and partition coefficients,
Sims et al. [1999] showed that there has to be at least an
order of magnitude decrease in upwelling velocity from
the tholeiitic systems to the alkali basalt systems. These
spatial variations in upwelling are broadly consistent with
predictions from geodynamic models for a plume-like
upwelling beneath Hawaii.
5.3.5. Influence of Lithospheric Mantle
[103] Melt generation in within-plate tectonic settings is
not necessarily restricted to upwelling asthenospheric
mantle, and melting within the overlying lithosphere is
also possible in both oceanic and continental environments. Hydrous phases such as amphibole and phlogopite
can be present at intermediate pressures within the oceanic
and continental lithosphere following metasomatic enrichment episodes, and their presence is known both from
xenolith samples and also indirectly from trace element
signatures in certain magmas. These minerals will preferentially incorporate Ba (and thus Ra too) relative to Th,
and if present at the few percent level in a mantle source,
this will lead to reduced (226Ra/230Th) or even 226Ra
deficits in magmas (e.g., Grande Comore [Claude-Ivanaj
et al., 1998], Samoa [Bourdon and Sims, 2003], and the
Azores [Claude-Ivanaj et al., 2001]).
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[104] In continental within-plate settings the lithospheric
mantle is likely to have a more significant role in influencing the compositions of the erupted magmas. Large lateral
variations in thickness will potentially limit the extent to
which the underlying asthenospheric mantle can decompress and melt, if at all. Metasomatic fluid and melt enrichments of the continental lithosphere will produce significant
mineralogical and compositional heterogeneities, and its
long stabilization history (1 – 3 Ga) means that trace element
enrichments will potentially develop significant radiogenic
isotopic compositions distinct from the asthenosphere
[Hawkesworth et al., 1990]. Trace element and radiogenic
isotope data can therefore be used to assess whether a
particular sample was likely to have been derived from
the asthenosphere or from enriched metasomatized lithospheric mantle.
[105] This approach has been used for within-plate magmas from the western United States [e.g., Perry et al., 1987;
Kempton et al., 1991]. Asmerom [1999] and Asmerom et al.
[2000] found that low– 143Nd/144Nd lavas (inferred to be
lithospheric melts) had 230Th-238U equilibrium, whereas
high – 143Nd/144Nd lavas (inferred to be asthenospheric
melts) had large 230Th excesses of 10– 40% (Figure 20a),
but both types had significant 231Pa excesses (Figure 20b).
The large 231Pa excesses rule out the possibility that the
230
Th-238U equilibrium values in the lithospheric melts are
simply the result of radioactive decay during slow magma
transport and long crustal residence. Modeling indicated
that the combination of 231Pa excesses and 230Th-238U
equilibrium require a spinel peridotite source, consistent
with the relatively shallow depth to the base of the lithosphere around the margins of the Colorado Plateau where the
lithospheric melts are found. It should be noted that equilibrium 230Th-238U values are not a distinguishing characteristic of all lithosphere-derived melts, as low-143Nd/144Nd
samples from other western United States locations can
have 230Th excesses (Figure 20a). A good example are
lavas from SW Utah [Reid and Ramos, 1996], where the
lithosphere is at least 100 km thick [Wang et al., 2002],
thick enough for garnet peridotite to be stable in the lower
20 or so km, thus enabling 230Th excesses to be generated.
It is also possible, though, for samples with relatively low
143
Nd/144Nd to have 230Th excesses simply as a result of
mixing between a lithospheric melt with even lower
143
Nd/144Nd and 230Th-238U equilibrium (generated in the
spinel-peridotite stability field) and an asthenospheric melt
with a large 230Th excess.
[106] Highly potassic magmas are generally interpreted as
melts of metasomatized regions of the lithospheric mantle.
Data from three examples (Tibet [Cooper et al., 2002],
Wudalianchi, China [Zou et al., 2003], and Gaussberg,
Antarctica [Williams et al., 1992]) all show very low
143
Nd/144Nd and significant 230Th excesses of between 7
and 60% (Figure 20a). For the Gaussberg lamproites and
the Wudalianchi potassic basalts the small within-suite
variations in 230Th-238U disequilibria correlate with radiogenic isotope compositions indicating a role for source
heterogeneity, likely to be the result of variable extents of
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these basalts were derived by 5– 7% melting of a slow
upwelling mantle source, consistent with evidence for
some lithospheric extension in the area, although very
small degree melting (0.3 – 0.5%) of a static source was
also permissible (i.e., using a ‘‘time-independent’’ melting
model) and highlights the ambiguities still present in such
modeling. Thus the degree of U series disequilibria in
continental within-plate lavas depends on many factors,
including the tectonic setting (thinned rift zone versus
thick lithosphere), the location of melting (asthenosphere
versus lithosphere), the source mineralogy (garnet-bearing
versus spinel-bearing mantle, which reflects the depth of
melting), and source compositional heterogeneities.
Figure 20. (a) (230Th/238U) versus 143Nd/144Nd for
within-plate lava suites: Gaussberg lamproites, Antarctica
[Williams et al., 1992], Tibet trachyandesites [Cooper et
al., 2002], Wudalianchi potassic basalts, China [Zou et al.,
2003], and Auckland alkali basalts, New Zealand [Huang
et al., 1997]. For the continental basalts from the western
United States, open circles are lithospheric melts from
Reid [1995] and Reid and Ramos [1996], while the data
from Asmerom and Edwards [1995], Asmerom [1999], and
Asmerom et al. [2000] are divided into lithospheric melts
(shaded circles) and asthenospheric melts (solid circles).
(b) (230Th/238U) versus (231Pa/235U), showing the distinction between asthenosphere-derived (high 143Nd/144Nd)
magmas with large 230Th excesses (solid circles) and
lithosphere-derived (low 143Nd/144Nd) magmas with small
or zero 230Th excesses (shaded circles) from the western
United States [Asmerom, 1999; Asmerom et al., 2000].
Field for oceanic island basalts (OIB) and MORB samples
is shown for reference (data sources as for Figure 6).
metasomatism of a garnet peridotite source. Major and
trace element data on the Wudalianchi potassic basalts
indicate that the source is a phlogopite-bearing garnet
peridotite, and in this region the lithosphere is known to
be 120 km. Zou et al. [2003] preferred a model in which
5.4. Subduction Zone Magmatism
5.4.1. Background to Melt Generation
at Subduction Zones
[107] Studies of elemental and isotopic variations in
subduction zone lavas over the last few decades, combined
with constraints from experimental petrology, have lead to a
broad consensus as to the dominant source components
involved in the genesis of subduction zone magmas (see
reviews by Hawkesworth et al. [1993] and Pearce and
Peate [1995]). Melt generation takes place predominantly in
the mantle wedge that overlies the subducting slab, as a
result of the lowering of the mantle peridotite solidus by
addition of water-rich fluids. These fluids are released by
dehydration reactions in the slab, primarily from the subducting basaltic crust but potentially from serpentinized
mantle as well. Many subduction zone magmas contain a
significant additional contribution from subducted sediments, and there is some evidence that the sediment
component may be added to the mantle wedge as a melt
before the fluid addition event that leads to arc magma
generation [e.g., Reagan et al., 1994; Elliott et al., 1997;
Turner and Hawkesworth, 1997]. The resulting magmas
may subsequently interact to some extent with the overlying
crust during ascent to the surface [e.g., Davidson, 1996].
[108] Subduction zones probably represent the most complex melting environment and, unlike at mid-ocean ridges,
the primary controls on melting are poorly established. Melt
generation will be influenced by the amount of fluid
introduced from the downgoing slab, the thermal structure
of the mantle wedge, and decompression in the mantle
wedge. Significant compositional variations both along
individual arcs and globally between arcs are also partly
controlled by variations in the amount and lithology of
sediments being subducted [e.g., Plank and Langmuir,
1993] and in the extent of depletion of the mantle wedge
due to prior melt extraction in the rear arc region [e.g.,
Woodhead et al., 1993]. A recent study by England et al.
[2004] argues that the location of arc volcanoes is critically
dependent on the thermal structure of the mantle wedge
rather than on the depth to the subducting slab.
[109] Fluids clearly play a major role in generating
subduction zone magmas, and it is likely that they will
have a significant influence on the extent and sense of
disequilibria between U series nuclides in these magmas.
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Figure 21. (a) The 238U-230Th equiline diagram for historic Mariana arc lavas [Elliott et al., 1997] plus
sketched model interpretation. (b) The 238U-230Th equiline diagram for Tonga arc lavas (shaded symbols)
and Vala Fa back arc lavas (open symbols) [Turner et al., 1997a; Peate et al., 2001b]. Arc samples with
low SiO2 (<55 wt %) and the Valu Fa back arc samples all have similar 143Nd/144Nd, indicating a similar
mantle source but with variable addition of a U-rich fluid, and they define an ‘‘isochron’’ age of 50 ka.
(c) (230Th/238U) versus Ba/Th for Tonga and Marianas arc lavas. (d) (226Ra/230Th) versus Ba/Th for
Tonga and Marianas arc lavas.
This is because of the contrasting chemical behavior of the
different U series nuclides in the presence of fluids. U and
Ra should behave like large ion lithophile elements (e.g.,
Ba, Rb, K, and Sr) and be readily mobilized in oxidized
aqueous-rich fluids, whereas Th and Pa should behave like
the relatively immobile high field strength elements (e.g.,
Zr, Nb, and Ti) [e.g., Brenan et al., 1995; Keppler, 1996].
Thus there is great potential to use 238U-230Th, 226Ra-230Th,
and 231Pa-235U disequilibria to provide constraints on the
timing of fluid addition to the mantle wedge source beneath
arcs. Turner et al. [2000c, 2003a] provide comprehensive
recent reviews of the origin and interpretation of U series
disequilibria in the subduction zone environment. A critical
aspect in the interpretation of U series disequilibria data
in subduction-related rocks is distinguishing the relative
contributions of elemental fractionation caused by fluid
addition from that caused by the partial melting process.
5.4.2. The 238U Excesses in Subduction Zone Magmas
and Links to Slab Fluid Addition
[110] A key feature of the subduction zone environment is that many lavas have significant 238U excesses
[e.g., Newman et al., 1984; Gill and Williams, 1990;
Hawkesworth et al., 1997a], and many of these lavas also
have much higher U/Th ratios than MORB or within-plate
lavas (Figure 7a). Observations of 238U-excess in other
tectonic settings are very rare and only of small magnitude,
with the exception of carbonatites (Figures 6 and 7). In
subduction zone lavas the highest 238U excesses are found
in samples with the lowest Th contents [e.g., Condomines
and Sigmarsson, 1993]. These samples are mainly from
trace element-depleted intraoceanic arcs such as Tonga and
the Marianas (Figures 21a and 21b), and they also tend to
have the strongest trace element evidence for addition of a
slab-derived fluid (e.g., high Ba/Th (Figure 21c)). The
mobility of uranium in fluids is governed by the redox
environment, as it is only fluid mobile under oxidizing
conditions where it can exist as U6+ rather than U4+,
unlike Th which only exists in the relatively immobile
Th4+ state. Compositional data on subduction zone peridotite samples show that conditions in the subarc mantle
are markedly more oxidizing than those in oceanic or
ancient cratonic mantle, indicating that the infiltrating slabderived fluids are oxidizing enough for U6+ to be the
dominant uranium species [e.g., Parkinson and Arculus,
1999]. Thus the simplest explanation of the elevated U/Th
and Ba/Th ratios in many subduction zone lavas is the
addition of a U-Ba-rich slab fluid to the mantle wedge
source. The observation that many subduction zone samples
have large 238U excesses indicates that this fluid addition
was, on average, a relatively recent event (<380 ka).
Along-arc correlations between plate convergence rate
and (230Th/238U) in the Alaska-Aleutian arc further suggest
that 238U excesses are linked to the relative size of the
fluid flux [George et al., 2003].
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[111] Lavas erupted at active spreading centers situated
behind the main volcanic arc front within back arc basins
show a diverse range of compositions. They include samples showing significant input from slab-derived fluids as
well as samples virtually indistinguishable from normal
mid-ocean ridge basalts. Thus these back arc lavas provide
a means to assess the relative influences of mid-ocean-ridge
style decompression melting processes and addition of
slab-derived fluids on 230Th-238U disequilibria. In the
two examples that have been studied in detail (Lau Basin
[Peate et al., 2001b] and Scotia Basin [Fretzdorff et al.,
2003]), lavas with 230Th excesses and lavas with 238U
excesses have both been found (Figure 6). In both cases
the samples with the 238U excesses show the clearest trace
element evidence for enrichment in a slab-derived fluid
component, and they tend to be found closest to the arc
front. For the Lau Basin the transition from samples with
238
U excesses to those with 230Th excesses takes place at
about 250 km behind the trench. The Lau Basin samples
with 230Th excesses and trace element compositions
broadly similar to mid-ocean-ridge basalts still have
enriched water contents relative to typical mid-ocean-ridge
basalts.
[112] Protactinium should behave similarly to high-fieldstrength elements such as Zr and Nb and be relatively
immobile in slab-derived fluids relative to uranium. Thus,
by analogy with the 230Th-238U system, excesses of 235U
over 231Pa should be expected in typical arc lavas. It is
somewhat surprising therefore that most subduction zone
lavas in fact have (231Pa/235U) > 1 (Figures 6 and 7) [Pickett
and Murrell, 1997; Bourdon et al., 1999b; Thomas et al.,
2002; Dosseto et al., 2003; Regelous et al., 2003]. In
contrast to fluid addition, partial melting should generate
231
Pa excesses because DPa < DU for mantle sources. The
observation of (231Pa/235U) > 1 for most subduction zone
lavas suggests that partial melting has a greater effect than
fluid addition in controlling the sense and magnitude of
231
Pa-235U disequilibria. The only known exceptions are
some extremely trace element-depleted lavas from the
Tonga arc. Bourdon et al. [1999b] argued that the highly
depleted mantle wedge in this region (because of prior melt
extraction in the back arc [Ewart and Hawkesworth, 1987])
contains only a small proportion of clinopyroxene and
that this makes it difficult to fractionate Pa from U during
melting.
5.4.3. The 230Th Excesses in Subduction Zone
Magmas and Links to Melting Processes
[113] It is notable that a significant proportion of subduction zone magmas (20%) have 230Th excesses similar in
magnitude to values observed in many mid-ocean ridge and
within-plate lavas (Figures 6 and 7). These samples tend to
be found in particular tectonic settings within subduction
zones, such as in rear arc regions (e.g., Kamchatka and
Sunda) and areas of thick crust (e.g., Central America,
Alaska, and the Andes), and to be associated with intra-arc
rifts (e.g., Kamchatka and Vanuatu). The process responsible for producing these 230Th excesses is not necessarily
the same in each case, but it likely involves some melting-
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related fractionation of Th and U. The three main options
are the following.
5.4.3.1. Decompression/Dynamic Melting
[114] Excesses of 230Th are an almost ubiquitous signature of decompression melting in mid-ocean ridge and
within-plate settings because of ingrowth of 230Th as melt
and matrix migrate through the melting region at different
velocities (see section 5.1 and Figure 15). Various lines of
evidence have been used to suggest that a component of
decompression melting can be important in the generation
of arc magmas [Pearce and Peate, 1995] (e.g., global
interarc correlations of major elements with lithospheric
thickness [Plank and Langmuir, 1988], theoretical modeling
of melting of hydrated peridotite [Hirschmann et al., 1999],
and the existence of low-H2O basalts in some arc front
volcanoes [Sisson and Bronto, 1998]). Decompression
melting is likely to be responsible for the 230Th excesses
observed in some lavas from extensional environments such
as back arc basins (e.g., Lau Basin [Peate et al., 2001b] and
Scotia Basin [Fretzdorff et al., 2003]) and some intra-arc
rifts (e.g., Aoba in the Vanuatu arc [Turner et al., 1999]).
Turner et al. [2003a] argued that the association of 230Th
excesses with lavas erupted through thick continental crust
rather than thin oceanic crust was consistent with models in
which the thickness of the overlying lithosphere influences
the mantle dynamics and partial melting within the wedge
[e.g., Plank and Langmuir, 1988], and George et al. [2003]
suggested that dynamic melting processes produced the
230
Th excesses in certain continental Alaskan lavas. Thomas
et al. [2002] showed that 230Th excesses can also be
produced by small degrees of melting using a flux melting
model, in which the extent of fluid addition drives the
melting process.
5.4.3.2. Slab Melting
[115] Thermal modeling indicates that melting of the
subducting slab is possible in the specific case where the
subducting oceanic crust is young and hot [Peacock et al.,
1994]. At depths >65 km the basaltic crust will be converted
to eclogite (garnet + clinopyroxene), and so slab melts
should be generated in the presence of residual garnet. This
will give rise to melts with 230Th excesses and trace element
features such as high La/Yb and Sr/Y. Addition of a slabderived melt has been suggested to explain the 230Th
excesses and trace element features of lavas in the austral
Andes in southern Chile [Sigmarsson et al., 1998b] and in
the central Kamchatkan depression [Dosseto et al., 2003].
5.4.3.3. Crustal Assimilation
[116] Bourdon et al. [2000b] showed that crustal assimilation can, at least in the case of Parinacota volcano in
Chile, lead to 230Th excesses if the crustal material is added
as a melt from a garnet-bearing residue. This will potentially
only affect lavas passing through very thick continental
crust as garnet is only stable at depths >30 km.
5.4.4. The 230Th/232Th Variations in Subduction Zone
Lavas: Influence of Subducted Sediments
[117] The Th budget of subduction zone lavas is largely
dominated by the addition of sedimentary material because
mantle rocks have much lower Th contents than continental-
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Figure 22. (a) The 238U-230Th equiline diagram for
samples from the Sunda arc [Turner and Foden, 2001;
Turner et al., 2003b], Kamchatka arc [Turner et al., 1998;
Dosseto et al., 2003], and the Nicaraguan part of the
Central American arc [Thomas et al., 2002]. (b) The
238
U-230Th equiline diagram for lavas from Mijakejima
volcano, Izu arc [Yokoyama et al., 2003].
derived sediments [e.g., Plank and Langmuir, 1993; Elliott
et al., 1997; Hawkesworth et al., 1997a]. Simple mass
balance calculations show that addition of just 1% of
typical subducted sediment [Plank and Langmuir, 1998]
will contribute 80% – 98% of the Th in a mantle wedge
source [e.g., Hawkesworth et al., 1997b]. Thus the Th
isotope composition of a mantle wedge source will be
largely governed by the 230Th/232Th of the added sediment, which can be estimated directly from the U/Th ratio
of the subducted sediment, assuming that the sediment is
in secular equilibrium and that it is added by bulk mixing.
However, there are several lines of evidence that suggest
that the sediment material might be transferred as a melt,
which can change the U/Th value of the added sedimentary component. Hawkesworth et al. [1997a] showed that
arc lavas globally display a broad negative correlation
between Th/Ce and 143Nd/144Nd consistent with variable
addition of sediment to a MORB-like mantle wedge.
However, samples with the highest sediment contribution
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(i.e., those with the lowest 143Nd/144Nd) had Th/Ce values
higher than most estimated subducted sediment compositions, which suggested that this elemental ratio had been
modified by partial melting of the sediment. In the
Marianas arc the samples with the highest sediment
contribution lie near the equiline on Figure 21a, with
(230Th/232Th) 1.0, and yet the bulk composition of the
subducting sediments is estimated to have (230Th/232Th)
0.6. Elliott et al. [1997] argued that the (238U/232Th)
value of the sediment was modified from 0.6 to 1.0 by
partial melting but that this sediment melt was then added
to the mantle wedge at least 380 kyr prior to the fluid
addition event so that the sediment-modified wedge had
time to ingrow 230Th so as to be close to secular
equilibrium (see Figure 21a).
[118] Samples from sediment-dominated arcs such as the
Sunda arc (as inferred from trace element and radiogenic
isotope data) tend to have low U/Th and 238U-230Th
disequilibria values close to the equiline (Figure 22a). This
is because addition of a slab-derived fluid has little
influence on 238U-230Th disequilibria because of the high
Th and U contents of the sediment-modified mantle wedge
[e.g., Hawkesworth et al., 1997a]. Most subducted sediments, with the exception of carbonates and U-rich hemipelagic deposits, have U/Th values lower than average
MORB mantle. Lavas from some subduction zones, notably
parts of Central America (Nicaragua) and Kamchatka have
high (230Th/232Th) values of 1.7– 2.6, greater than observed
in any MORB samples (Figure 22a). For Kamchatka the
preferred model is for an old (>380 ka) fluid enrichment
episode, which produced high U/Th ratios that through
subsequent radioactive decay has lead to the observed
high (230Th/232Th) values [Turner et al., 1998; Dosseto
et al., 2003]. However, in the case of Nicaragua, high
(230Th/232Th) values are consistent with the addition of
sedimentary material with high U/Th (and thus high
(230Th/232Th)) together with some effect also from an old
fluid-enrichment episode [McDermott and Hawkesworth,
1991; Reagan et al., 1994; Thomas et al., 2002].
5.4.5. Systematic 238U-230Th Disequilibria Variations
Within Individual Island Arcs
[119] In detail, samples from individual subduction zones,
mainly those from trace element depleted island arcs, often
define broad linear arrays on the 238U-230Th equiline
diagram. If these trends are interpreted as isochrons, they
have apparent ages ranging from 10 to 200 kyr prior to
eruption (summarized byTurner et al. [2000c]). The clearest
example is shown by data from the oceanic Marianas arc
that define an age of 30 ka (Figure 21a) [Elliott et al., 1997].
Data from the Tonga arc and back arc system show a more
scattered trend (Figure 21b) with a slope corresponding to
an age of 50 ka [Turner et al., 1997a; Peate et al., 2001b].
There has been considerable debate as to the age significance of these pseudo isochrons, and some of the various
interpretations are sketched on Figure 21a.
[120] Earlier studies tended to interpret the slope of the
array as an isochron, for which the age would represent
the time elapsed since a discrete episode of fluid addition
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to the mantle wedge source. This interpretation requires
several assumptions: (1) The fluid contains no Th, so
variable fluid addition initially produces a horizontal array
on the equiline diagram; (2) the mantle wedge has the
same 230Th/232Th throughout the arc; and (3) melting
processes do not modify the 238U-230Th systematics.
Another possible interpretation for the arrays is that the
fluid does contain some Th, so the initial fluid addition
may produce a sloped rather than a horizontal array.
Similarly, if the mantle wedge has variable U/Th, addition
of a constant flux of fluid containing just U will have a
greater effect on the more depleted parts of the wedge with
higher initial U/Th, thereby producing a sloped data array.
Several recent studies have also argued that melting
processes beneath arcs can produce sloped arrays on the
238
U-230Th equiline diagram [e.g., Elliott et al., 2001;
Thomas et al., 2002; Bourdon et al., 2003a; George et
al., 2003], which will be discussed further in section 5.4.7.
[121] It is important to realize that most studies have been
regional in coverage with only one or two samples from
individual volcanoes within an arc. Yokoyama et al. [2003]
have shown, through a detailed study of a single arc volcano
(Miyakejima volcano, Izu arc), that the 238U-230Th systematics can be locally very complex. The earliest lavas erupted
between 25 and 10 ka define a linear trend with an
apparent age of 22 ka, and the lavas erupted between 7
and 4 ka define a different linear trend with an apparent age
of 12 ka (Figure 22b). Both lines intersect the equiline at a
common (230Th/232Th) value of 1.3, which, together with
the uniform 143Nd/144Nd ratios in these samples, suggests
that the mantle wedge beneath Miyakejima was compositionally homogeneous prior to fluid addition. The more
differentiated samples erupted in the last few thousand years
define an array that is steeper than the equiline and
indicative of magma mixing, consistent with petrological
and other compositional data. Yokoyama et al. [2003]
argued that it was difficult to explain the two trends in the
least differentiated lavas by a flux-melting-type model [e.g.,
Thomas et al., 2002], and instead, they explained the data in
terms of the episodic events of fluid addition triggering melt
generation in the mantle wedge on a timescale of several
thousands of years.
5.4.6. The 226Ra-230Th Disequilibria and Implications
for Magma Ascent Rates
[122] Many subduction zone lavas have significant
226
Ra excesses (Figure 7c) [e.g., Gill and Williams,
1990; Turner et al., 2001a], and the question is whether
these 226Ra excesses result from melting processes or fluid
addition. The highest (226Ra/230Th) values are generally
found in the least differentiated lavas, indicating that the
226
Ra-230Th disequilibrium has a mantle origin. Turner
and Hawkesworth [1997] suggested that the 226 Ra
excesses were generated by partial melting and in that
way decoupled from the 230Th-238U disequilibrium caused
by addition of a slab fluid. However, the degree of
226
Ra- 230Th disequilibrium in some arc lavas, with
226
( Ra/230Th) up to 7, is far greater than in MORB and
OIB, which generally have (226Ra/230Th) < 3, and gener-
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ating such extreme values by decompression melting alone
would require extremely small porosities (<0.0001%
[Turner et al., 2001a]). Furthermore, within individual
arcs, 226Ra excesses generally show a reasonable correlation with Ba/Th (Figure 21d), which suggests a link with
addition of a slab-derived fluid [Turner et al., 2001a].
Bourdon et al. [2003a] also noted that the positive
correlation of (226Ra/230Th) with Sr/Th shown by samples
from both the Kamchatka arc and the Tonga-Kermadec
cannot be due to melting because Sr is more compatible
than Th during mantle melting (unlike for Ba, which is
more incompatible than Th). Instead, the trends can be
explained by the fluid mobility of Sr and Ra relative to
Th. A further link between a slab-derived signature and
226
Ra excess is provided by the positive correlation
between (226Ra/230Th) and 10Be/9Be in lavas from southern Chile [Sigmarsson et al., 2002], as 10Be is unambiguously derived from subducting sediments.
[123] The principal 230Th-238U and 226Ra-230Th fractionation appears to occur when the oxidizing fluid is released
from the dehydrating slab. We are therefore faced with an
apparent discrepancy between the timescales for fluid
transfer as indicated by these two disequilibria pairs. The
arrays on the 230Th-238U equiline diagram can be interpreted
as the result of a fluid addition event 10– 200 kyr ago, but if
226
Ra excesses were produced during the same event, they
would have already decayed back to equilibrium. The
observed 226Ra excesses therefore suggest a very recent
fluid addition to the mantle wedge within the last few
thousand years. If this interpretation of the 226Ra excesses
is correct, that they were generated as the fluids were
released from the subducted slab, then this has significant
implications for the mechanism of fluid transfer and ascent
rate of magmas. The fluid has to be transferred from the slab
to the melting zone within the mantle wedge, and then the
resulting melt has to be transferred from 80 to 100 km depth
to be erupted at the Earth’s surface within just a few
thousand years, at most, of the fluid release event. This
implies rapid fluid transfer, probably via hydraulic fracturing [Davies, 1999], rapid melt segregation and ascent by
channeled flow, followed by minimal residence in magma
chambers within the crust [Clark et al., 1998; Turner et al.,
2000b; Turner et al., 2001a; Sigmarsson et al., 2002;
Yokoyama et al., 2003].
[124] Feineman and DePaolo [2003] argue that minerals
in the mantle wedge might not be in secular equilibrium for
226
Ra-230Th. The 226Ra generated by decay of 230Th can
readily diffuse out of clinopyroxene and potentially into
phlogopite if present, as Ra is compatible in phlogopite. If
the added slab fluid is stabilized in the wedge by growth
of phlogopite, then as the mantle flow brings this material
to the melting zone over a finite time (perhaps tens of
kiloyears), the 238U excesses would decrease because of
radioactive decay, while the phlogopite would maintain a
steady state elevated (226Ra/230Th) by diffusive reequilibration. Feineman and DePaolo [2003] argue that this
model lessens the constraints for a rapid fluid transfer from
the slab as previously suggested based on 226Ra/230Th
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[e.g., Turner et al., 2001a], although it still implies rapid
melt ascent rates (1 km yr1) from the initial site of
melting through the mantle wedge and crust to the surface.
[125] The question remains as to how the 230Th-238U and
226
Ra-230Th data can be reconciled. Several studies have
shown reasonable correlations between (230Th/238U) and
(226Ra/230Th) for particular subduction zone suites [Reagan
et al., 1994; Chabaux et al., 1999; Sigmarsson et al., 2002],
and these authors suggested that both U and Ra were added
at the same time to the wedge by a single slab fluid less than
a few thousand years ago, such that the arrays on the U-Th
equiline diagrams are essentially mixing trends between the
fluid and the wedge compositions. However, Bourdon et al.
[2003a] have argued that this would imply unrealistically
high fluid partitioning values for Th and hence that a
significant proportion of the 230Th in arc lavas must come
from ingrowth from excess 238U over a time period >10 kyr.
Experimental work by Schmidt and Poli [1998] indicates
that slab dehydration will probably result in the multistep or
continuous release of fluids over a significant range of
depths. Turner et al. [2000b] modeled a simple two-stage
addition approximation in an attempt to explain data from
the Tonga arc. The initial fluid addition will remove virtually
all the 238U and 226Ra but leave 230Th and 232Th behind in
the slab. Further dehydration of the slab will release liquids
that contain 226Ra produced from decay of the 230Th
remaining in the slab but no U. The 226Ra excesses produced
by the first fluid addition event will have decayed away by
the time of the subsequent fluid additions, so Turner et al.
[2000b] argued that the 238U excesses reflected the timing of
the initial fluid addition, whereas the 226Ra excesses
reflected the last increment of fluid addition.
5.4.7. Subduction Zone Melt Generation Models
[126] Compared to mid-ocean ridges and even withinplate setting, the development of realistic quantitative melting models for subduction zones is still in its infancy,
largely because of the complexity of the melt generation
process. There are uncertainties about the details of the
mechanisms that must be incorporated into a melting model
(e.g., dynamics of mantle flow in wedge, how the fluid is
transferred from the slab and how it migrates within the
wedge to the melting zone, and dynamic melting versus
fluid-fluxed melting), and many of the critical parameters
are poorly constrained (e.g., thermal structure of mantle
wedge and element partitioning in fluids). Nevertheless,
initial exploratory attempts are being made to model details
of melting in arcs as constrained by U series disequilibria
[Thomas et al., 2002; Bourdon et al., 2003a; George et al.,
2003; Yokoyama et al., 2003].
[127] Bourdon et al. [2003a] argued that equilibrium
porous flow melting models [e.g., Spiegelman and Elliott,
1993], in which there is extensive reequilibration between
melt and residual mantle throughout the melting column, are
unlikely in the arc setting for several reasons: (1) The fast
melt migration times inferred from the 226Ra data leave little
time for melt equilibration. (2) The elevated 187Os/188Os
isotope compositions of some arc lavas, if confirmed as a
slab signature, would not survive if there was significant
RG1003
melt equilibration with the mantle wedge. (3) Explaining the
(231Pa/235U) and (226Ra/230Th) data of arc lavas by melting
of a fluid-modified source using a equilibrium porous flow
model requires very high melting rates, and by implication
very fast upwelling velocities (200 cm yr1), compared to
geologically more reasonable rates using a dynamic melting
model. Therefore initial models have investigated the consequences of dynamic melting [e.g., McKenzie, 1985;
Williams and Gill, 1989] taking place either subsequent to
fluid addition [Elliott et al., 2001; Bourdon et al., 2003a;
George et al., 2003] or contemporaneous with fluid addition
(fluxed ingrowth melting [Thomas et al., 2002; Bourdon et
al., 2003a]), and these two scenarios are illustrated in
Figure 23.
5.4.7.1. Dynamic Melting Following Fluid Addition
[128] Bourdon et al. [2003a] considered the situation in
which fluid addition and partial melting are separate events,
perhaps with a finite time interval between them. Assuming
reasonable values for partition coefficients, porosity, and
fluid-induced disequilibria, the modeling achieved values
for (231Pa/235U) and (226Ra/230Th) typical of arc lavas for
melting rates on the order of 103 – 104 kg m3 yr1. At
larger melting rates, there is not sufficient time to ingrow
231
Pa in the residual mantle, and at smaller melting rates the
initial 226Ra excess from fluid addition will have decayed
away. The dynamic melting process will also produce linear
arrays on the U-Th equiline diagram reminiscent of an
isochron [see also Elliott et al., 2001; George et al.,
2003]. Slower melting rates produce steeper arrays (and
hence older ‘‘isochron ages’’) than faster melting rates, and
the slopes are also influenced by the choice of partition
coefficients for U and Th that will reflect the mineral mode
and extent of depletion of the mantle wedge.
5.4.7.2. Fluxed Ingrowth Melting
[129] In the model of Thomas et al. [2002], partial
melting occurs as an immediate response to continued fluid
addition. Fresh hot mantle is dragged through the zone of
slab fluid addition because of corner flow in the mantle
wedge, and it is assumed that the transfer of fluid from the
slab, the partial melting, and the melt transport are all
essentially instantaneous. Flow of the solid mantle, on the
other hand, takes a long enough time (105 – 106 years) to
pass through the melting region to allow ingrowth of 231Pa
and 230Th from U remaining in the residues of melt
extraction. Arc magmas are then formed by integrating
melts from different parcels of mantle that have experienced
different amounts of fluid addition and different extents of
melting. This model also produces linear arrays on the U-Th
equiline diagram, as a consequence of mixing of melts that
have experienced fluid addition and melting over a range of
time intervals.
[130] Although these two models are physically quite
different, parameters can be chosen such that either model
can generally fit the observed U series disequilibria in
magmas from different subduction zones [Bourdon et al.,
2003a]. A critical issue remains the extent of mobility of Th
and Pa in fluids, and other constraints may come from a
better understanding of the thermal and dynamical regime
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Figure 23. Sketch cross sections to illustrate the principal features of two melting models for
subduction zone magmas: (a) dynamic melting following fluid addition and (b) fluxed ingrowth melting
(Figures 23a and 23b taken from Turner et al. [2003a], with permission from Mineralogical Society of
America.) (c and d) Predictions for U series disequilibria variations for model shown in Figure 23a as
various parameters such as melting rate are varied (assumes DPa = 0.0001, DU = 0.003, DTh = 0.002, and
DRa = 0.00001) (Figures 23c and 23d taken from Bourdon et al. [2003a]). In Figure 23c the solid lines
represent models for different melting rates, at a constant porosity of 3%, while the dashed curve is a
model for a melt depleted mantle wedge with DU = 0.0003 and DTh = 0.00026. In Figure 23d the curves
illustrate how (226Ra/230Th) and (231Pa/235U) are predicted to vary as a function of melting rate for two
different matrix porosity values: Solid curves assume an initial (231Pa/235U) = 0.8 (i.e., prior addition of a
U-bearing fluid), while the dashed curves assume an initial (231Pa/235U) = 0.8. (e and f) Predictions for U
series disequilibria variations for model shown in Figure 23b as various parameters are varied (assumes
DPa = 0.000005, DU = 0.0051, DTh = 0.00204, DRa = 0.00005, melting rate = 0.35 g m3 yr1, and
critical porosity = 1%) (Figures 23e and 23f taken from Thomas et al. [2002], with permission from
Elsevier). Shaded lines denote melting of an enriched mantle (with U = 0.05 ppm and Th = 0.157 ppm).
Solid lines denote melting of a depleted mantle (with U = 0.03 ppm and Th = 0.078 ppm); tick marks
indicate the average extent of melting in percent.
within the mantle wedge as the two models also differ in the
timescales of melting and the size of the inferred melting
region.
5.5. Final Comments on U Series Disequilibria
and Melting Processes
[131] Interpretation of U series disequilibria data on
mantle-derived magmas has been one of the driving forces
leading the development of more realistic models for melt
generation and transport in the Earth. Observed variations
of U series disequilibria in lavas clearly preserve important
information about several physical parameters of the melt
generation and transport process (e.g., melting rate, porosity, solid upwelling rate, melt ascent rate, and extent of
melt-solid equilibrium during melt ascent). However, any
quantitative interpretation of U series disequilibrium data
is highly dependent on the exact choice of melting model,
as well as the absolute values for partition coefficients
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[e.g., Spiegelman and Elliott, 1993; Condomines and
Sigmarsson, 2000; Lundstrom, 2003; Bourdon and Sims,
2003; Turner et al., 2003a]. Furthermore, the role for
diffusional processes is still uncertain [e.g., Saal and Van
Orman, 2004; Feineman and DePaolo, 2003, 2004;
George et al., 2004]. Several authors [e.g., Lundstrom et
al., 1998a; McKenzie, 2000] have stressed that it still
remains to be seen whether the existing melting models give
a complete picture of the principal physical processes
involved in melt generation and movement. In order for
the full potential of U series disequilibria data to be realized,
independent geochemical, geophysical, and fluid dynamical
investigations are required to add additional constraints on
realistic melting models for different tectonic environments.
For geochemists an important step forward is the collection
of comprehensive analyses on well-dated samples that
include major and trace element analyses, radiogenic and
stable isotope data, and volatile analyses (e.g., H2O, CO2, Cl,
and SO2) in addition to (230Th/238U), (226Ra/230Th), and
(231Pa/235U).
[132] Despite these reservations, U series disequilibria
data have provided several robust observations about melting processes. For example, melting beneath mid-ocean
ridges initiates in the region of the garnet-spinel transition
or deeper, and the melts can be extracted at very low
porosities of 1% or less. At subduction zones most lavas
preserve evidence for a recent (380 kyr) addition of a
U-rich fluid derived from the subducted slab, and the
simplest interpretation of the (226Ra/230Th) disequilibria
data is that melt ascent rates through the mantle wedge
and crust are extremely rapid (1 km yr1 on average).
6.
SUMMARY AND FUTURE DIRECTIONS
[133] Recent U series disequilibria studies of igneous
rocks have indicated how different age information can be
obtained from the crystals and the liquid in the same rock
and how the time spent by the magma and the residue in the
melt generation zone are very different. In a number of
cases, phenocryst minerals have been shown to have formed
after the differentiation processes responsible for the whole
rock compositions. Concordant ages have long been
regarded as a test of the geological significance of radiometric ages, but recent studies have highlighted the additional information that may be available in models that seek
to reconcile different ages obtained on the same material.
[134] The progress of the last few decades has highlighted
a number of aspects that require further consideration and
study, and these include the following.
[135] 1. For all global tectonic settings, there is a need to
move from regional surveys to detailed local studies, using
sample suites that have good stratigraphic and age control
and that are well characterized in terms of complete
elemental and radiogenic isotope data. These should
include 238U-234U-230Th-226Ra and 235U-231Pa data and
should be situated where there are independent geophysical constraints on crustal and mantle structure and geodynamic parameters.
RG1003
[136] 2. The same arguments hold on the much smaller
scale of whole rocks and minerals, in that it is important to
understand fully the textural relations of the phases being
analyzed. This is increasingly possible as in situ analytical
techniques for both trace elements and isotopes become
more sensitive. The goals are then to model systems that
yield discordant age information, to integrate age information from radiogenic isotopes with that from crystal size
distribution and elemental diffusion profiles, to contrast the
record from cumulates and lavas, and to explore the extent
to which preeruption processes can be linked with events in
the volcanic stratigraphic record. Do earlier differentiation
or crystal growth events revealed through isotope disequilibria data coincide with dated events from the eruptive
record of the volcanic system or changes in the style of
volcanic behavior?
[137] 3. We need to develop physically realistic models
consistent with constraints from compositional data, experimental petrology, and thermal evolution and geophysics
and also to explore how the textures of igneous rocks reflect
the conditions and timescales of crystallization.
[138] 4. More experiments are required to improve our
knowledge of element partitioning as a function of pressure,
temperature, and volatile contents, in particular at small
degrees of melting. Melting rates as a function of source
composition and water content, the role of diffusion during
melting and melt percolation, and the effects of grain
boundary partitioning need to be better constrained.
[139] 5. For subduction zone magmas, determining the
solubilities of U series nuclides in plausible slab-derived
fluids, especially for Th and Pa, remains a high priority.
These would then be combined with improved physical
models of slab dehydration reactions, melt generation and
transport, and the thermal and convective regime within the
mantle wedge.
[140] ACKNOWLEDGMENTS. Funding for this study was
provided by the Danish National Research Foundation through a
grant to the Danish Lithosphere Centre and by NERC and the
Leverhulme Trust at Bristol. Discussions with Steve Blake, Jon
Blundy, Tim Elliott, Rhiannon George, Dan Morgan, Mark
Reagan, Marcel Regelous, Simon Turner, and Georg Zellmer have
all contributed to the ideas summarized here. Comments from Chris
Hieronymus and Julia Shaw on an earlier version of this manuscript
are gratefully acknowledged. Julia Shaw is also thanked for
allowing us to include unpublished Pa data from the Reykjanes
Ridge and SWIR on Figures 6 and 7. Ingrid Ukstins Peate is
thanked for invaluable help in drafting some of the figures. We also
thank editor Tom Torgersen and three reviewers for their comments
that improved the presentation of ideas in this review.
[141] The Editor responsible for this paper was Thomas
Torgersen. He thanks an anonymous cross-disciplinary reviewer
and Michael Murrell and another anonymous technical reviewer.
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C. J. Hawkesworth, Department of Earth Sciences, University of
Bristol, Wills Memorial Building, Queens Road, Bristol BS8 1RJ, UK.
D. W. Peate, Department of Geoscience, University of Iowa, 121
Trowbridge Hall, Iowa City, IA 52242, USA. ([email protected])
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