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Chapter1
General Introduction
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1-1. The Earth’s continental crust
The outermost part of the solid Earth is the crust, the base of which is defined by the
Mohorovicic discontinuity.
The crust can be classified into two crustal types: oceanic crust
comprising about 60 % of the crust by area and continental crust comprising the other (Cogley,
1984).
The oceanic crust is mafic and about 1–11 km thick.
In contrast, the continental crust is
from felsic to mafic and about 30–40 km thick (e.g., Taylor and McLennan, 1989; Rudnick and
Fountain, 1995).
The presence of two different types of crusts results in a unique topographic
expression of a bimodal frequency of elevation in the Earth, whereas Venus has only one
predominant elevation.
The moon has also a mafic oceanic crust and a thick (up to 75 km) felsic
continental crust, and exhibits a bimodal topographic expression.
However, there are significant
differences between the moon’s and Earth’s continental crusts in the composition and the
mechanism of formation.
The moon’s continental crust is anorthositic and was formed by
solidification of a magma ocean at its birth (Warren, 1985).
On the other hand, the Earth’s felsic
continental crust is typically granitoid and has been formed by magma extraction from the mantle
and/or mafic crust (e.g., Taylor and White, 1965; Anderson, 1982; Shirey and Hanson, 1984;
Martin, 1986; Kushiro, 1990).
Consequently, the continental crust renders Earth unique among all
known planets (Taylor, 1989), and is probably linked to the presence of liquid water on Earth
(Campbell and Taylor, 1985).
In addition, the continental crust is a major reservoir of the
incompatible trace elements in the silicate Earth, such as Rb, Ba, K, P, Pb, Th, and U (e.g., Rudnick
and Fountain, 1995), and has a great influence on Earth’s surface environment, such as the
compositions of the seawater and the atmosphere, and the surface temperature (e.g., Urey, 1953;
Walker et al., 1981).
Therefore, knowledge of the growth history of the continental crust is
essential to understanding the evolution of our planet to its present form.
To discuss the growth of the continental crust, first we should know the characteristics of
the Earth’s continental crust.
Because lower part of continental crust is relatively inaccessible,
our view on the lower continental crust comes largely from a few uplifted slices of the crust in
collisional orogens and from xenoliths in younger volcanics.
Petrological and geochemical
studies on the rock samples with seismic data demonstrated that the lower continental crust is
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composed of rocks in the granulite facies and that it is lithologically heterogeneous with mafic
average composition (Rudnick and Taylor, 1987; Christensen and Mooney, 1995; Rudnick and
Fountain, 1995).
On the other hand, because the upper continental crust is accessible to geological and
geochemical studies, it is reasonably well constrained. The upper continental crust is composed of
granitoids with minor mafic rocks, and its average composition is similar to granodiorite (Condie,
1993).
The granitoids can be subdivided into four groups according to their composition with
source rock characteristics (Table 1-1); S-type (sedimentary source type), I-type (igneous source
type), M-type (mantle source type), and A-type (anorogenic type) (Chappell and White, 1974;
White and Chappell, 1977; Loisell and Wones; 1979; White, 1979).
S-type granitoids are
characterized by low concentrations of Na2O and CaO, and low ratios of (Na2O + K2O +
CaO)/Al2O3 and Na2O/K2O.
The granites typically exhibit radiogenic Sr isotopic compositions,
and sometimes contain xenoliths of sedimentary rocks.
These observations indicate that the
source of S-type granitoids contains sedimentary rocks (Chappell and White, 1974; White and
Chappell, 1977).
I-type granitoids are characterized by high concentrations of Na2O and CaO, and
high ratios of (Na2O + K2O + CaO)/Al2O3 and Na2O/K2O.
often found in the granitoids.
unradiogenic.
Xenoliths of mafic igneous rocks are
In addition, their Sr isotopic compositions are relatively
For these reasons, the source of I-type granitoids is considered to be mafic igneous
rocks (Chappell and White, 1974; White and Chappell, 1977).
M-type granitoids are
characterized by markedly high ratios of Na2O/K2O and CaO/(Na2O + K2O).
compositions are significantly unradiogenic.
Their Sr isotopic
A-type granitoids are distinguished from others by
low contents of Al2O3 and low ratios of (Na2O + K2O)/CaO and (MgO + CaO)/(FeO + Fe2O3).
Further details of the classification of granitic rocks are well summarized in Barbarin (1999) and
Frost et al. (2001).
Although each type of the granitoids was formed in various regions and generations, its
mode has changed with time.
predominant.
In Phanerozoic, the formation of S-type granitoids was relatively
In Proterozoic, A-type granitoids constitute a significant fraction of the granitoids.
On the other hand, common granitoids formed in Archean, termed as Archean TTG (tonalitetrondhjemite-granodiorite; Jahn et al., 1981), belong to I-type or M-type granitoids, and are also
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characterized by high Sr/Y and La/Yb ratios with lack of negative Eu anomaly (e.g., Martin, 1986).
The change of the mode of formed granitoids with time can be interpreted as reflecting the
change of the dominant formation process of granitoid with time, whereas there are various
granitoid formation models for each generations and types (e.g., Shirey and Hanson, 1984; Martin,
1986; Kushiro, 1990; Rudnick, 1995; Smithies, 2000).
As an example, Martin (1986) suggested
that the dominant formation process in Archean time is partial melting of hydrated young and warm
slab, leaving a residue of garnet-amphibolite or eclogite, and that it after Archean, in contrast, is
partial melting of the hydrated mantle wedge.
However, alternative formation process for
Archean TTG, such as partial melting of hydrous basaltic material previously underplated beneath
thickened crust rather than subducted oceanic crust, were also proposed (e.g., Rudnick, 1995;
Smithies, 2000).
These controversies on the formation processes of granitoids are long-standing
problems, and further constraints on them will provide new insights into the evolution of
continental crust and mantle.
1-2. Continental growth models
So far, many continental growth models have been presented based on various methods
and standpoints (Fig. 1-1).
These models illustrate the various factors that may have contributed
to the continental growth.
Here we shall look at them, but firstly we must know the direct
evidence for the continental growth what we have at this moment.
Fig. 1-2 is a map showing distribution of orogenic belts of the world at the present.
The
oldest known terrestrial rocks are 4.03 Ga granitoids in the Acasta Gneiss Complex, northwestern
Canada (Stern and Bleeker, 1997; Bowring and Williams, 1999); no crustal rocks have been found
from the first five hundred million years of Earth’s history (Nutman et al., 2001).
On the other
hand, the oldest identified terrestrial materials (>4.06 Ga) are minerals of zircon with ages up to 4.4
Ga from the Yilgarn Craton in Western Australia that occur as detrital grains in ca. 3.0 Ga
metasediments (e.g., Froude et al., 1983; Compston and Pidgeon, 1986; Wilde et al., 2000; Wyche
et al., 2004) or as xenocrystic grains in ca. 2.6 Ga granitoids (Nelson et al., 2000).
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Mineralogical
and geochemical studies on the zircons indicate that some of them crystallized from granitoid
magma (Mojzsis et al., 2000; Cavosie et al., 2005; Crowley et al., 2005; Watson and Harrison,
2005).
This suggests the formation of continental crust only a few hundred million years after the
Earth’s formation.
It also should be kept mind that rocks older than ca. 3.6 Ga are very rare at the
present (e.g., Nutman et al., 2001).
A major issue in the study of the continental growth history is how to estimate the
difference of its “actual” growth rate (i.e. the net gain in mass of continental crust per unit of time)
from the “minimum” growth rate (the rightmost growth curve MR in Fig. 1-1) based on the present
age distribution of orogenic belts (Maruyama, 1995; Maruyama and Liu, 1998).
The difference
can be caused by three recycling processes: (1) the erosion of crustal materials and subsequent
deposition and diagenesis (intra-crustal recycling; Tugarinov and Bibikova, 1976; Rino et al., 2004),
(2) formation of granitoids from remelting of older continental crustal materials (crustal reworking;
Hurley and Rand, 1969; Veizer and Jansen, 1979, 1985; Condie, 1998), and (3) subduction of
continental crustal materials into the mantle (crust-mantle recycling; Armstrong, 1968, 1981; Fyfe,
1978; Reymer and Schubert, 1984).
Below, the proposed continental growth models are
summarized primarily in the light of these three recycling processes.
1-2-1. Models considering intra-crustal recycling
Tugarinov and Bibikova (1976) preliminarily estimated the growth rate from the areal
distribution of age provinces of the continental basement rocks based on the whole-rock Rb-Sr and
zircon U-Pb age determinations.
However, there are large areas in which no information on the
age of the basement rocks can be obtained, because they are overlain by sedimentary rock covers.
Hence, there is a controversy on whether the data from restricted area correspond to the age
distribution of the mean of continental crust.
Recently, Rino et al. (2004; RN in Fig. 1-1) have determined U-Pb ages of ca. 1500
detrital zircons from large rivers in North and South America, in order to determine more precise
growth rate.
Zircon is ubiquitous accessory minerals in continental crust and has a high resistance
to intra-crustal recycling and metamorphic events through geological time.
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In addition, detrital
zircons in the river mouths originally derived from the (hidden) basement continental crust in the
extensive drainage area.
Hence U-Pb age distributions of zircons from the rivers reflect the
continental growth rates of the continental crust in their source regions.
Based on the obtained
results, they demonstrated that the continental crust was episodically formed at 2.8–2.5 Ga and
2.3–1.9 Ga.
However, it should be noted that these results do not provide conclusive information
on whether the formed crust is juvenile or reworked: we cannot know whether continental crust
grown at the time or not.
1-2-2. Models considering intra-crustal recycling and crustal reworking
Hurley and Rand (1969; H & R in Fig. 1-1) calculated the continental growth rate by
using the present age distribution of basement continental crust that was estimated from whole-rock
Rb-Sr and K-Ar age determinations, and rough crustal reworking rates in each generations that was
estimated based on the difference of mean initial Sr isotope compositions of various rock samples
from that of average Earth (Fig. 1-3; see also Section 1-2-4).
be applied to large part of the continental crust.
However, this method also could not
In addition, note that in spite of their
consideration of both intra-crustal recycling and crustal reworking, the curve is clearly righter than
that of Rino et al. (2004), and almost overlaps with the “minimum” growth curve.
This can be
because of the isotopic disturbance, especially injection of radiogenic Sr from a fluid, on the
whole-rock Rb-Sr and K-Ar isotope systems during later metamorphic events, although they tried
to correct empirically for the effect of isotopic disturbance (only isotopic reset) by later thermal
events.
Veizer and Jansen (1979; V & J in Fig. 1-1) also calculated the growth rate in a similar
manner by using the present age distribution reported by Tugarinov and Bibikova (1976) and rough
crustal reworking rates which was estimated from the difference of the mean initial Sr isotope
compositions of various granites from that of oceanic basalts.
However, they, as well as Hurley
and Rand (1969), could not estimate a single precise reworking rate because the rate most likely
varies with time and between regions.
McCulloch and Bennett
(1994; M & T in Fig. 1-1) estimated the growth curve primarily
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based on Nd chondritic (CHUR) model ages (Fig. 1-3; see also Section 1-2-4) of granitoids and
sedimentary rocks from North American, North American, and Scandinavian cratons. Along with
the results, they suggested episodic formation of juvenile continental crust at 3.6–3.5 Ga, 2.7–2.6
Ga, and 2.0–1.8 Ga.
However, since the data are obtained from restricted area of some cratons, it
is doubtful whether the estimated growth rate corresponds to that of the mean of the continental
crust.
1-2-3. Models considering crust-mantle recycling and constancy of freeboard
Armstrong firstly emphasized the significance of crust-mantle recycling with the growth
model that the volume of continental crust has remained constant since ca. 3.0 Ga (i.e., no-growth
model; Armstrong, 1968, 1981; AM in Fig. 1-1).
The idea was derived from the geological
observations that suggest a near-steady state of average thickness of continental crust since ca. 3.0
Ga (Condie, 1973) and the near constancy of freeboard (i.e., mean elevation of continental surface
relative to sea level) since the end of Archean (Wisen, 1972, 1974).
Under the assumption that
there are no changes in the volumes of the seawater and the depth of oceanic basins with time, the
geological observations imply that the continent must have not grown from ca. 3.0 Ga.
If this is
the case, and in order to accommodate observed Pb, Sr, and Nd isotopic compositions of terrestrial
rocks (e.g., Moorbath et al., 1977; O’Nions et al., 1979; see also Section 1-2-4) to the model, then
the continental crust must have recycled into and well mixed with the mantle (Armstrong, 1968,
1981).
In contrast, Moorbath argued that low-density felsic crust cannot be lightly subducted into
the mantle, and therefore the continental crust must have grown throughout geological time (i.e.,
progressive-growth model).
Reymer and Schubert (1984; R & S in Fig. 1-1) estimated the continental growth rate at
the present by using rates of continental addition (e.g., arc magmatisms and incorporation of
ophiolites) and of continental subtraction (mainly subduction of pelagic sediments), and arrived at a
net growth rate of ca. 1 km3/yr.
If the estimated growth rate is constant since the end of Archean,
Archean growth rates must have been 3–4 times the present rate on the average.
They also
claimed that the constancy freeboard demonstrated by Wisen (1972, 1974) requires a slight crustal
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growth with time, because the secular decline in the heat production of the mantle causes the
volume of seawater to increase and the ocean basins to deepen with time, and therefore that their
growth curve is consistent with the constancy of freeboard.
However, the assumption, that growth
rate has not changed since the Late Archean, is highly doubtful.
In addition, because of the large
uncertainties in the freeboard itself and in the controlling parameters such as atmospheric
composition, there is a great ambiguity in the growth models based on the constancy freeboard.
1-2-4. Models based on the isotope evolution of the depleted mantle
So far, Sr and Nd isotopic evolution curve for the depleted mantle have been used in
estimating growth rates (e.g., Allègre, 1982; O’Nions et al., 1979; AL and O’N in Fig. 1-1,
respectively).
The method is primarily based on the characteristic that Sm and Nd as well as Rb
and Sr are fractionated during the magmatic processes due to the difference of the incompatibility
between parent (Sm, Rb) and daughter (Nd, Sr) isotopes.
Hence extraction of continental crust
(enriched crust) with relatively low Sm/Nd and high Rb/Sr results in a complementary high Sm/Nd
and low Rb/Sr of the mantle (depleted mantle), and accordingly leads to the evolution of radiogenic
Nd with unradiogenic Sr in the depleted mantle and correspondingly unradiogenic Nd with
radiogenic Sr in the continental crust, as compared to bulk silicate earth.
It should be noted here
that we could evaluate whether a formed crust is juvenile or reworked by comparing its initial (at
the formation) Sr and/or Nd isotope compositions with those of the bulk silicate earth or the
depleted mantle at the time (see Section 1-2-2).
Because initial Sr and Nd isotopic compositions of oceanic basalt correspond to the
isotopic compositions of the mantle at the time, we can estimate the depleted mantle isotopic
evolution curve by determining the initial isotopic compositions of oceanic basalts with different
ages.
Figs. 1-3a and b are the Nd and Sr isotopic evolution curves for the depleted mantle,
illustrating the reasonable derivation the depleted mantle curves from those of bulk silicate earth
through geological time.
Allègre (1982) deduced from the evolution curves for the depleted
mantle that continental crust has grown throughout geological time, but mainly during Precambrian.
However, the growth models based on the depleted mantle isotopic evolution (e.g., Allègre, 1982;
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O’Nions et al., 1979) are highly controversial mainly because of the uncertainty of the size of the
depleted mantle contributing to the continental growth, whereas the method could be useful to
qualitatively evaluate whether an enriched crust existed or not at specific time, especially in
Hadean.
1-2-5. Models based on the secular changes in compositions of sedimentary rocks
McLennan and Taylor (1982; M & T in Fig. 1-1) demonstrated that post-Archean clastic
sediments differ from Archean clastic sediments by the presence of negative Eu anomaly, and that
there is no significant change in the magnitude of Eu anomaly within post-Archean clastic
sedimentary rocks.
Mass balance modeling of REE abundances in the Archean and post-Archean
sedimentary rocks for the observations suggests that the minimum ratio of post-Archean to Archean
upper continental crust is about 4 : 1, corresponding to the continental growth of ca. 80 % of the
present volume by 2.5 Ga.
However, because they compared largely greenstone sediments in the
Archean with cratonic sediments in the post-Archean, it is difficult to test their conclusion.
In
addition, Condie (1993) argued that not only post-Archean but also Archean upper crust have
negative Eu anomaly.
Veizer (1989) tried to determine the secular change in the Sr isotopic composition of
seawater using carbonate samples, because the isotopic composition is controlled by the input from
eroded continental crust (relatively high Sr isotopic composition) and from hydrothermal water
(mantle-like isotopic composition).
The Sr isotopic compositions obtained from the carbonates
suggest that seawater Sr isotopic composition followed the mantle Sr isotopic evolution curve until
ca. 2.5 Ga, and increased rapidly at ca. 2.5–2.0 Ga, and eventually increased steadily thereafter.
Based on the alleged increase at ca. 2.5 Ga, Veizer (1989) proposed the rapid continental growth at
the time.
However, Gibbs et al. (1986) pointed out the possibility that the alleged change resulted
from comparing Archean greenstone carbonate with post-Archean cratonic carbonate.
Subsequently, Condie (1992) demonstrated that there is no significant secular change in the relative
Sr isotopic compositions from the mantle within each tectonic setting sample, providing no positive
evidence for the rapid continental growth at ca. 2.5 Ga.
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1-3. This study
As described above, many researchers have evaluated the effect of one or more of the
three recycling processes and proposed various continental growth curves depending on their
consideration and discipline.
history.
We can learn much from them for the study of continental growth
Especially two lessons are taken here; one is the utility of zircon due to its ubiquitous
presence in the continental crust and robustness to intra-crustal recycling and metamorphic events
throughout geological time (see Section 1-2-1), and the other one is the versatility of Sm-Nd and/or
Rb-Sr isotope systematics, which can provide information not only on the degree of geochemical
differentiation event in the silicate Earth (e.g., continental crust formation) but also on the mode of
continental crust formation (juvenile crust formation vs. crustal reworking; see Sections 1-2-2 and
1-2-4).
On the other hand, several problems to be resolved still remain.
Firstly, the degree of
crustal reworking has to be estimated more quantitatively, whereas McCulloch and coworkers (e.g.,
McCulloch and Wasserburg, 1978; McCulloch and Bennett, 1994; see also Section 1-2-2) obtained
overview of the age spectra of the juvenile continental crust formation.
Secondly, a long-standing
question of whether no-growth since the late Archean (e.g., Fyfe, 1978; Armstrong, 1981; Reymer
and Schubert, 1984) or progressive-growth through geological time (e.g., Moorbath, 1979), more
specifically a subject concerning the importance of crust-mantle recycling (see Section 1-2-3), must
be investigated further.
The controversy originated mainly from the lack of our knowledge on the
nature of early continental crust, such as its volume and composition, due to the extreme rarity of
terrestrial materials before 4.0 Ga.
One way to deduce the volume of the early continental crust is
to estimating the degree of differentiation of early silicate Earth based on isotope systematics (e.g.,
Sm-Nd, Lu-Hf) of early Archean rocks, although it is not a conclusive information (see Section 12-4).
In order to provide more rigid constraints on its volume and composition, it is important to
discover and study additional very old terrestrial materials.
The purpose of this study is two-fold: (1) to quantitatively and precisely determine the
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rates of crustal reworking and juvenile continental formation, and (2) to constrain on the volume
and composition of early continental crust.
We have approached to the objectives primarily using
U-Pb ages and Lu-Hf isotopic compositions of zircons.
U-Pb zircon geochronology is one of the
most precise and utilized dating methods to determine crystallization ages of their host igneous
rocks.
On the other hand, though it has been widely recognized that the Lu-Hf isotope system of
zircon can be used as a tracer to understand not only the time-integrated differentiation of the
silicate Earth but also the mode of formation of its host igneous rock, in analogy with the Sm-Nd
isotope system (see details in Fig. 1-3 and Section 1-2-4), its application had been retarded.
This
is because Hf is an analytically unfavorable element for conventional thermal ionization mass
spectrometry (TIMS) due to high ionization energy of Hf.
The limitation is now eliminated
because Hf isotopes can be measured with high precision and accuracy by new multiple collectorinductively coupled plasma mass spectrometry (MC-ICPMS) technique (e.g., Halliday et al., 1998).
For the first purpose, we determined U-Pb ages and initial Hf isotopic compositions of ca.
400 detrital zircon grains from the Mississippi River sand.
Whereas the distribution of the U-Pb
ages reflects a specific sequence of major magmatic events in the source region of the zircons (as
was pointed out by Rino et al., 2004), that of the Hf model ages, calculated using the U-Pb ages and
initial Hf isotopic compositions, reflects the rate of juvenile continental crust formation.
In
addition, by comparing these two age distributions, we can determine precise crustal reworking
rates.
The obtained results and discussion are described in Chapter 3.
For the second purpose, we carried out works on the Acasta Gneiss Complex in
northwestern Canada, that contains the oldest known rocks with ages as old as 3.9–4.0 Ga (e.g.,
Bowring et al., 1989).
The works consist of detailed geology, U-Pb zircon geochronology and
zircon Lu-Hf isotope geochemistry.
In Chapter 4, based on the geological and geochronological
data, we argue the tectonothermal history of the Acasta Gneiss Complex.
At the same time, we
also report the discovery of a very old zircon xenocryst with an age of 4.2 Ga within an Acasta
gneiss.
In Chapter 5, the obtained initial Hf isotopic compositions of the Acasta gneisses and its
implications for the nature of the early continental crust are discussed.
Most of the U-Pb and Lu-Hf isotopic data discussed in this study have been provided by
means of laser ablation-inductively coupled plasma mass spectrometry (LA-ICPMS).
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Prior to
analyze the zircon samples, we have improved the analytical method of in-situ zircon Lu-Hf
isotope analysis in order to obtain precise and accurate data from small ablation pits (~35 µm).
Chapter 2 is the description of this achievement.
Finally, conclusions of this study are summarized in Chapter 6.
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Figure 1-1
Representative of continental growth models proposed in the literature; MR, Maruyama (1995);
H&R, Hurley & Rand (1969); RN, Rino et al. (2004); V&J, Veizer & Jansen (1979); M&B,
McCulluch & Bennett (1994); M&T, McLennan & Taylor; O’N, O’Nions et al. (1979); AL, Allégre
(1982); B, Brown (1979); AM, Armstrong (1981); R&S, Reymer & Schubert (1984); F, Fyfe
(1978).
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Figure 1-2
Map showing orogenic belts of the world (modified after Maruyama, 2002).
The locations of the
Acasta Gneiss Complex in the Slave Craton, northwestern Canada, and the Yilgarn Craton, Western
Australia are also represented.
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Figure 1-3
Nd (a) and Sr (b) isotopic evolution curves for the bulk silicate earth (BSE) and depleted mantle
(DM) (modified after Allègre, 1982).
schematically shown.
The evolution curve of the continental crust is also
In the case that a rock was formed by remelting of older continental crust
(crustal reworking) at T, its initial Nd and Sr isotope ratios are low and high, respectively, relative
to that of the depleted mantle at T.
In addition, if the Sm/Nd ratio of the reworked continental
crust is known, we can estimate the formation age of the reworked continental crust (i.e., juvenile
continental crust) by calculating its model ages (TCHUR and TDM are chondritic and depleted mantle
model ages, respectively).
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Table 1-1
Average composition (wt%) for each type of granite (Whalen et al., 1987)
I-type
S-type
A-type
M-type
SiO2
69.17
70.27
73.81
67.24
TiO2
0.43
0.48
0.26
0.49
Al2O3
14.33
14.10
12.40
15.18
Fe2O3
1.04
0.56
1.24
1.94
FeO
2.29
2.87
1.58
2.35
MnO
0.07
0.06
0.06
0.11
MgO
1.42
1.42
0.20
1.73
CaO
3.20
2.03
0.75
4.27
Na2O
3.13
2.41
4.07
3.97
K2O
3.40
3.96
4.65
1.26
P2O5
0.11
0.15
0.04
0.06
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