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ORIGINAL PAPER
J. W. Valley ÆJ. S. Lackey ÆA. J. Cavosie
C. C. Clechenko ÆM. J. Spicuzza ÆM. A. S. Basei
I. N. Bindeman ÆV. P. Ferreira ÆA. N. Sial ÆE. M. King
W. H. Peck ÆA. K. Sinha ÆC. S. Wei
4.4 billion years of crustal maturation: oxygen isotope ratios
of magmatic zircon
Received: 7 April 2005 / Accepted: 11 July 2005 / Published online: 9 November 2005
Springer-Verlag 2005
Abstract Analysis of d
18
O in igneous zircons of known
age traces the evolution of intracrustal recycling and
crust-mantle interaction through time. This record is
especially sensitive because oxygen isotope ratios of
igneous rocks are strongly affected by incorporation of
supracrustal materials into melts, which commonly have
d
18
O values higher than in primitive mantle magmas.
This study summarizes data for d
18
O in zircons that
have been analyzed from 1,200 dated rocks ranging
over 96% of the age of Earth. Uniformly primitive to
mildly evolved magmatic d
18
O values are found from
the first half of Earth history, but much more varied
values are seen for younger magmas. The similarity of
values throughout the Archean, and comparison to the
composition of the ‘‘modern’’ mantle indicate that d
18
O
of primitive mantle melts have remained constant
(±0.2&) for the past 4.4 billion years. The range and
variability of d
18
O in all Archean zircon samples is
subdued (d
18
O(Zrc)=5–7.5&) ranging from values in
high temperature equilibrium with the mantle (5.3±
0.3&) to slightly higher, more evolved compositions
(6.5–7.5&) including samples from: the Jack Hills (4.4–
3.3 Ga), the Beartooth Mountains (4.0–2.9 Ga), Bar-
berton (3.5–2.7 Ga), the Superior and Slave Provinces
(3.0 to 2.7 Ga), and the Lewisian (2.7 Ga). No zircons
from the Archean have been analyzed with magmatic
d
18
O above 7.5&. The mildly evolved, higher Archean
values (6.5–7.5&) are interpreted to result from ex-
change of protoliths with surface waters at low tem-
perature followed by melting or contamination to create
mildly elevated magmas that host the zircons. During
the Proterozoic, the range of d
18
O(Zrc) and the highest
values gradually increased in a secular change that
documents maturation of the crust. After 1.5 Ga, high
d
18
O zircons (8 to >10&) became common in many
Proterozoic and Phanerozoic terranes reflecting
d
18
O(whole rock) values from 9 to over 12&. The ap-
pearance of high d
18
O magmas on Earth reflects non-
uniformitarian changes in the composition of sediments,
and rate and style of recycling of surface-derived ma-
terial into magmas within the crust.
Electronic Supplementary Material Supplementary material is
available for this article at http://dx.doi.org/10.1007/s00410-005-
0025-8
Communicated by J. Hoefs
J. W. Valley (&)ÆC. C. Clechenko ÆM. J. Spicuzza
J. S. Lackey ÆA. J. Cavosie
Department of Geology, University of Wisconsin, Madison,
WI 53706, USA
E-mail: valley@geology.wisc.edu
Tel.: +1-608-2635659
Fax: +1-608-2620693
M. A. S. Basei
Department de Mineralogia Geotectonica,
University Sao Paulo, Sao Paulo, Brazil
I. N. Bindeman
Department of Geological Sciences, University of Oregon,
Eugene, OR 97403, USA
V. P. Ferreira ÆA. N. Sial
NEG-LABISE, Department of Geology,
Federal University of Pernambuco, Recife, PE 50670-000, Brazil
E. M. King
Department of Geography and Geology, Illinois State University,
Normal, IL 61790, USA
W. H. Peck
Department of Geology, Colgate University, Hamilton,
NY 13346, USA
A. K. Sinha
Virginia Polytechnic Inst., Blacksburg, VA USA
C. S. Wei
School of Earth and Space Sciences,
University of Science and Technology of China,
Hefei, Anhui, 230026, China
Contrib Mineral Petrol (2005) 150: 561–580
DOI 10.1007/s00410-005-0025-8
Introduction
Oxygen isotopes in zircon
Zircon is a common accessory mineral in igneous rocks
and preserves the most reliable record of both magmatic
oxygen isotope ratio (d
18
O, Valley 2003) and magmatic
age (U-Th-Pb, Hanchar and Hoskin 2003). Several fac-
tors combine in zircon to create a robust and retentive
geochemical record, including: high temperatures of
mineral stability and melting, slow diffusion rates for
cations and anions, chemical inertness, and hardness.
High contrast cathodoluminescence and other imaging
techniques distinguish domains of growth zoning from
igneous and subsolidus overgrowth, resorption, and ra-
diation damage. While many common minerals are
readily altered by metamorphic, hydrothermal, or di-
agenetic processes, zircons are generally not affected.
Zircons with heavy radiation damage or postmagmatic
alteration can be identified and avoided prior to analy-
sis. No other mineral permits d
18
O(magma) to be cou-
pled to age of crystallization with such confidence.
Oxygen isotope ratios of magmas reflect the d
18
Oof
magmatic source rocks and contaminants. With rare
exceptions, the mantle is a remarkably homogeneous
oxygen isotope reservoir (Eiler 2001) and igneous zir-
cons in high temperature equilibrium with mantle mag-
mas have average d
18
O = 5.3±0.3&(1 SD, Valley
et al. 1998). Even small deviations from the mantle value
of d
18
O are readily apparent. Fractional crystallization
can result in higher whole rock (WR) values of d
18
Oby
up to 1&in more silicic magmas, however the value of
d
18
O(Zrc) remains approximately constant because the
fractionation, D
18
O(WR-Zrc), increases at nearly the
same rate as d
18
O(WR) due to the greater abundance of
higher d
18
O minerals, e.g., quartz and feldspar, in the
evolving, more silicic magmas. The change in d
18
O(WR)
is increased if temperature decreases significantly during
differentiation, however the effects of variable tempera-
ture on d
18
O(Zrc) are minor due to small intermineral
fractionations at magmatic temperatures and because
zircon fractionations are intermediate among rock-
forming minerals (i.e., zircon is neither the highest nor
lowest d
18
O mineral in a rock, Valley 2003; Valley et al.
2003). Therefore, significant deviations of d
18
O(Zrc)
from the mantle value are the direct or indirect result of
intra-crustal recycling, i.e., magma interaction with su-
pracrustal materials that ultimately derived their evolved
d
18
O from low temperature processes on or near the
surface of the Earth where oxygen isotope fractionations
are large.
Oxygen isotope reservoirs
The d
18
O values of common crustal materials are sum-
marized in Fig. 1.Bothd
18
O(WR) and d
18
O(Zrc) are
shown. The fractionation, D
18
O(Zrc-WR), varies with
mineralogy and can be approximated as a linear func-
tion of wt % SiO
2
for igneous rocks at magmatic tem-
peratures. Values of D
18
O(Zrc-WR) vary from 0.5&
for mafic rocks to 2&for granites according to the
relation:
D18OðZrc WRÞ¼d18 OðZrcÞd18 OðWRÞ
0:0612ðwt:%SiO2Þþ2:5
(Valley et al. 1994; Lackey 2005). For comparison with
the crust, two vertical lines show the mantle range of
d
18
O(Zrc) at 5.3±0.3&. Fresh basalts (WR) are close,
but slightly above the range for mantle zircon, but altered
basalts plot at higher or lower values depending on the
temperatures of interaction with surface waters. Like-
wise, in ophiolite sequences, low d
18
O gabbros have been
altered by high temperature hydrothermal fluids while
the high d
18
O basalts were altered at low temperatures
(Gregory and Taylor 1981; Eiler 2001). The general-
ization that low temperature water–rock interactions
cause high d
18
O also applies for continental and oceanic
sediments that uniformly plot at much higher values re-
flecting interaction with surface water. The range of ig-
neous zircons for various rock types is more subdued in
d
18
O, reflecting the magmatic values, and generally above
the mantle value. Low d
18
O magmas have been intensely
studied in a few localities, especially sub-volcanic en-
vironments, but are not a volumetrically significant
component of the crust (Balsley and Gregory 1998).
-10 0 10203040
δ18O ‰ VSMOW
Oceanic Crust
Seawater
Sediments (W.R.)
Siliceous oozes
Carbonate oozes
Pelagic clays
Clastic sediments
Igneous Rocks (W.R.)
Altered basalts
Fresh basalts
Layer-3 gabbros
Continental Crust
Meteoric water
Sediments (W.R.)
Sandstones
Limestones
Shales
Cherts
Igneous Rocks (W.R.)
I-type granites
S-type granites
Igneous Zircons
Mantle
Archean
Proterozoic
Phanerozoic
to -55
"Mantle Zircon" δ18O = 5.3 ± 0.3
Fig. 1 Typical values of d
18
O for sediments, igneous rocks, and
igneous zircons (modified from Eiler 2001). Ticks for continental
sediments represent average values for the Archean. The narrow
field at 5.3±0.3&represents d
18
O of zircons in high-temperature
equilibrium with the mantle (plotted at 1SD). Zircons from
primitive magmas fall near this field, and values above 6.5&result
from recycling of supracrustal material. The distribution of low
d
18
O zircons is uncertain before 150 Ma and are not shown (see
text)
562
Average values are shown for modern sandstones,
limestones, and shales in Fig. 1. The compositions of
Precambrian sedimentary rocks are lower than modern
sediments and have average d
18
O(WR) values of: 16.7&
for shales (Land and Lynch 1996); 20&for carbonates
(Shields and Veizer 2002), 13–14&for sandstones (Blatt
1987); and 24–28&for cherts (Blatt 1987; Perry and
Lefticariu 2003). Some chemical sediments are system-
atically lower in d
18
O as a function of increasing age
leading to provocative proposals of secular changes in
d
18
O of sediments and oceans through time (Walker and
Lohmann 1989; Burdett et al. 1990; Land and Lynch
1996; Muehlenbachs 1998; Wallmann 2001; Shields and
Veizer 2002; Perry and Lefticariu 2003; Knauth and
Lowe 2003; Veizer and Mackenzie 2003). Changes
through time in the composition or availability of sedi-
ments for magmatic recycling will influence the d
18
Oof
any resultant igneous rocks.
Evolved d
18
O in magmas
High values of magmatic d
18
O, above that derived from
the mantle, are most often found in granitic rocks and
attributed to melting or assimilation of sediments, al-
tered volcanics, or other supracrustal rocks of near-
surface genesis. We distinguish such ‘‘intra-crustal re-
cycling’’, where supracrustal materials are melted or
contaminate magma that intrudes continental crust,
from ‘‘mantle recycling’’, where continental crust is
subducted and returned to the mantle.
Many questions of granite genesis and the definition
of granite types are beyond the scope of this contribu-
tion. The d
18
O values of S- and I-type granites in Fig. 1
are characteristic of type localities in SE Australia and
other Phanerozoic examples (O’Neil and Chappell 1977;
O’Neil et al. 1977). As early as 3.1 Ga, granite plutons
are estimated to represent 20% of Archean exposure
(Condie 1993) including many that are peraluminous
(Sylvester 1994) and might contain a sedimentary com-
ponent. However, we avoid widespread application of
the S and I classifications (see Chappell and White 2001).
Magmatic values of d
18
O can also be shifted by as-
similation or remelting of altered igneous rocks. Mag-
matic cannibalization is common in plutonic complexes
where successive magmas intrude and may melt each
other. In this situation, early crystallized magmas are
often hydrothermally altered by water circulation pow-
ered by the heat of later magmas. The d
18
O (and dD) of
altered wall rock is shifted while other geochemical
systems are generally unaffected. Radiogenic isotope
systems cannot detect this process because of insufficient
time for ingrowth of daughter isotopes. Thus, some
melts that are partially or wholly produced in the crust
(from mantle-derived materials) may appear mantle-
derived. Analysis of d
18
O in zircons allows clear dis-
tinction of magmatic versus postmagmatic composition
and, in many instances, provides the only evidence for
cannibalization or wall rock contamination (Valley
2003). The distinction of first and second-generation
magmas is significant beyond the sphere of isotope
geochemistry, affecting estimates of the rate of heat and
mass transfer, and crustal growth.
This study reports oxygen isotope ratios for igneous
zircons with ages from 4.4 Ga to nearly the present. We
demonstrate the utility of zircon oxygen isotope ratios as
a monitor of magmatic chemistry, and highlight con-
trasting behavior between oxygen, the major element in
the crust and the mantle, and commonly applied trace
element and radiogenic isotope systems. One goal is to
test the generality of the observation that Archean
magmas in North America had uniform values of d
18
O
within 2&of the mantle (5.3&) while post-Archean
magmas were higher and more variable (Peck et al.
2000). A second goal is to examine the timing and causes
of this secular trend.
Techniques
Magmatic zircons of known age have been analyzed for
d
18
O from 1,200 rocks (Table 1, Appendix 1). For most
samples, U-Pb age was measured in previous studies by
thermal ionization mass-spectrometry (TIMS) or by ion
microprobe (SIMS). For a few samples, age was inferred
based on geochronology of associated rocks. All detrital
zircons (4.4–2.9 Ga) from the Jack Hills and the Bear-
tooth Mountains were analyzed in situ by ion microp-
robe for age and d
18
O. Whole rock chemical data are
available for some samples.
A majority of the d
18
O analyses were made at the
University of Wisconsin–Madison. Zircons separated
from igneous rocks were analyzed for d
18
O in samples
consisting of 1–2 mg, typically 100–1,000 zircons, that
were concentrated by standard crushing, gravimetric,
and magnetic procedures. For samples previously dated
by TIMS, aliquots of the original zircon separate were
obtained. Concordance of U-Pb ages provides an index
of radiation damage, and analysis of concordant sam-
ples enhances confidence in the reliability of d
18
O values
as primary (Valley 2003; Cavosie et al. 2005). In many
samples, more than one magnetic or size split was ana-
lyzed to test for variability and guard against significant
deviation of d
18
O due to inheritance of older cores. The
least magnetic zircons available were analyzed so as to
correspond as closely as possible to those that were
dated. In rare cases where detectable differences in d
18
O
are seen among different zircons from the same sample,
d
18
O for the least magnetic zircons is reported because
they display little or no evidence of radiation damage
(Valley et al. 1995). For zircon samples that were ori-
ginally separated for oxygen isotope or fission track
studies, age is typically reported from geochronology on
the same unit. Since 1999, zircons in the Wisconsin lab
have been soaked in concentrated HF at room tem-
perature for 8–12 h to dissolve impurities and metamict
material. Cold HF does not affect d
18
O of undamaged
zircons (King et al. 1998; Valley 2003). Clouded grains
563
Table 1 Age and oxygen isotope ratio of igneous zircons tabulated in Appendix 1, given as ESM, available at http://dx.doi.org/10.1007/
s00410-005-0025-8
Location Age d
18
O 1Std. # Dominant #
Range Zircon Dev. Rocks Lithologies Outliers
Ma Ave. permil d
18
O*
ARCHEAN
Jack Hills, Yilgarn craton, Australia 4404–3280 6.2 0.7 59 Detrital zircons
Beartooth Mountains, Wyoming province 3973–2936 6.2 0.5 10 Detrital zircons
Barberton, South Africa 3538–2740 5.53 0.67 11 Granite, tonalite
Superior Province Volcanics 2736–2691 5.57 0.48 45 Rhyolite, dacite 3
Superior Province Plutonic
Wabigoon Subprovince 3003–2680 5.65 0.52 36 TTG**, sanukitoid, gabbro
Quetico Subprovince 2688 6.83 1 Quartz-monzonite
English River Subprovince 2698–2697 6.69 0.21 3 TTG, sanukitoid, gabbro
Uchi Subprovince 2741–2700 5.90 0.34 5 TTG, sanukitoid
Wawa Subprovince 2728–2678 5.94 0.48 6 TTG, sanukitoid
Abitibi Subprovince 2720–2668 6.03 0.98 8 Syenite, monzonite, quartz-diorite
Lewisian 2700 5.48 0.46 2 Tonalitic orthogneiss
Slave Province 2694–2670 4.87 0.26 5 Tonalite
PROTEROZOIC
China 2560–2494 5.64 0.21 3 Granite, granodiorite
Brazil 2251–1894 5.33 0.71 16 Granitic-mafic orthogneiss
Trans Hudson 2597–1819 6.15 0.72 11 Granitic-tonalitic gneiss
Ukranian Shield 2695–1720 6.64 1.03 12 Granite, granodiorite, gabbro
Australia 1858–1806 7.11 0.79 6 Granite, granodiorite
Wisconsin 1860–1760 5.08 0.11 2 Granite, tonalite
Finland 1886–1573 6.94 0.95 23 Granite
Great Basin, western U.S. 1500 6.60 1.11 19 Pegmatite, orthogneiss
Laramie Anorthosite Complex 1340 7.35 0.21 5 Monzosyenite
Nain Anorthosite Complex 1330–1285 6.24 0.67 20 Granite, ferrodiorite, anorthosite
Grenville Province
Adirondack Mountains 1336–900 7.86 1.20 60 Granitic to mafic orthogneiss 1
Frontenac 1176–1160 11.34 1.63 8 Granitic to monzonitic orthogneisses
Quebec 1240–1077 7.64 1.41 13 Granitic orthogneiss, anorthosite
Grenville-age, Vermont 1154–1119 7.76 0.49 3 Augen gneiss
Grenville-age, Virginia and Maryland 1162–998 7.39 0.65 24 Augen gneiss, granitic orthogneiss
Virginia and Maryland (Neoproterozoic) 748–680 6.39 0.51 3 Granite
Uruguay (Various Ages) 2111–510 7.84 0.98 6 Granitic orthogneiss
Argentina (Various Ages) 1000–206 7.96 1.76 9 Granodiorite, orthogneiss
Brazil (Neoproterozoic)
Curitiba Microplate 590 6.20 0.28 2 Granite, diorite
NE Paran State 633–564 7.40 1.14 13 Granite to granodiorite
Paranagu and Monguagu Batholiths 620–567 7.20 0.98 10 Granite, tonalite
Serra do Mar Alkaline-Peralkaline Suite 604–540 5.58 0.79 21 Granite, rhyolite
Pien Batholith 618–605 6.02 0.55 7 Granite to diorite
Brusque Metamorphic Complex 638–580 7.25 0.35 2 Granite, syenite
Florianpolis Batholith 640–609 7.39 0.62 6 Granite
Pelotas Batholith 620–580 7.64 0.52 7 Granite
SE border of San Francisco Craton 653–632 7.03 1.05 2 Orthogneiss
Sao Rafael Pluton 627 5.98 0.17 9 Granite, quartz-monzonite
Emas Pluton 633 10.04 0.22 8 Granodiorite
Borborema province 880–581 8.73 0.92 34 Shoshonite,high-K calc-alkaline
Nubian shield, Israel 620 7.59 0.80 3
PHANEROZOIC
Grelo, Spain 310 6.17 1 Granodiorite
Greece 316–233 6.51 1.16 3 Granitic orthogneiss
Antarctica 183 6.09 1 Felsic dike
Western U. S.
Northern Sierra Nevada batholith 143–140 5.55 0.69 3 Granodiorite
Western Sierra Nevada batholith 145–93 6.77 0.82 36 Tonalite to gabbro
Eastern Sierra Nevada batholith 222–81 6.19 0.48 85 Granite to granodiorite
Central Sierra Nevada batholith 162–89 6.94 0.60 91 Granite to diorite
Southern Sierra Nevada batholith 117–81 7.82 0.72 45 Granite to diorite
Peraluminous Plutons, Sierra Nevada 165–86 7.95 0.46 35 Granite
Owens Valley/White Mountains 217–74 6.83 0.69 31 Granite to granodiorite
Death Valley 173 7.07 1 Monzonite
Idaho Batholith 100–45 6.97 0.78 29 Granite to granodiorite
564
were removed by hand picking and resistant zircons
were ground for analysis. Most zircon separates were
analyzed at least twice. Zircon powder is heated by CO
2
laser in a BrF
5
atmosphere to yield O
2
that is cryo-
genically purified, reacted with hot graphite, and ana-
lyzed as CO
2
in a dual-inlet gas-source mass-
spectrometer. Analyses are standardized by replicate
analyses (3 or more) on the same day of UWG-2 garnet
standard (d
18
O=5.8&) and reported in standard per mil
notation relative to VSMOW (Valley et al. 1995). Ty-
pical precision for these analyses is ±0.05&(1SD) and
accuracy relative to NBS-28 quartz standard is ±0.1&.
All of the published and unpublished data for
d
18
O(Zrc) that we are aware of are included in Appendix
1, except as noted in text. Approximately 60% of the
d
18
O(Zrc) data are previously published and another
20% are in manuscripts that are in preparation or review
(Appendix 1). The selection of samples was guided by
the goals of these previous studies and by the availability
of zircon concentrates. Thus the coverage is not perfectly
distributed through time and across all major geologic
terranes.
For detrital igneous zircon crystals (e.g., Jack Hills
metaconglomerate and Beartooth Mountains quartzite)
d
18
O and U-Pb isotopic age are correlated using in situ
analyses from the same crystal by ion microprobe. U-Pb
age was measured from 20 to 30 lm spots (<1 ng) by
SHRIMP II at Curtin University (Wilde et al. 2001;
Peck et al. 2001) or at the Chinese Academy of Geolo-
gical Sciences (Cavosie et al. 2004). For samples with
multiple spots, the oldest concordant age is reported.
Two generations of oxygen isotope data are reported. In
1999, d
18
O was measured by CAMECA ims-4f using a
10 kV beam of
133
Cs
+
, high energy offset (14.15 kV
total potential), and a single electron multiplier at the
University of Edinburgh (precision 1.0&, 1 SD, Peck
et al. 2001). In 2004, d
18
O was measured at the Uni-
versity of Edinburgh by CAMECA ims-1270 using high
mass-resolution and dual faraday detectors (precision
0.3&, 1 SD, Cavosie et al. 2005). A careful protocol for
analysis is documented in each study where standard
analyses bracket unknowns and represent 25–50% of all
analyses in each analytical session. The analytical pits
for oxygen isotope analyses made in 2004 are, when
possible, directly below the polished locations of age
measurements and were examined after analysis by
SEM. No artifacts of earlier analyses were detected, but
cracks and inclusions can be evaluated. These ion mi-
croprobe data represent sample sizes that are 10
6
times
smaller (ng vs. mg) than the laser fluorination data.
Most crystals are homogeneous in d
18
O within analytical
uncertainty and the average value is reported for single
crystals where multiple analyses were made.
Results
Archean
Zircon concentrates were analyzed for d
18
O from 121
Archean rocks from four geologic provinces on three
continents (Fig. 2, Table 1, Appendix 1). In addition, 67
detrital igneous zircons were analyzed by ion microp-
robe. The igneous Archean zircons have a small range of
Fig. 2 Map showing the exposed and inferred extent of Archean
rocks and areas studied (from Peck and Valley 2005)
Table 1 (Contd.)
Location Age d
18
O 1Std. # Dominant #
Range Zircon Dev. Rocks Lithologies Outliers
Ma Ave. permil d
18
O*
Great Basin 480–27 6.84 1.23 124 Granite
Timber Mountain/ Oasis Valley 12.8–11.3 5.75 1.07 9 Rhyolite, latite
Bishop Tuff 0.76 5.83 0.17 4 Rhyolite
Yellowstone 2–0.109 3.17 1.48 26 Rhyolite
China
Eastern China 126–98 4.62 0.57 37 A-type granite
Dabie Orogenic Belt 120 5.15 0.46 37 Post metamorphic granite
British Tertiary Igneous Province
Arran 58 6.48 0.97 3 Granite
Isle of Skye 58 3.23 1.48 21 Granite
Isle of Mull 58 4.25 1.44 3 Granite
*Not included in average.
**TTG tonalite, trondhjemite, granodiorite
565
d
18
O (5.0–7.4&). Taken together, the data yield an
average of d
18
O(Zrc) = 5.82 ±0.74&(Fig. 3a) and no
variability is seen with age or SiO
2
content throughout
the period 4.4–2.5 Ga (Figs. 4and 5). However, even
within this limited range of values, there are correlations
with the host rock type. Most samples have primitive
d
18
O(Zrc) values (5.5±0.7&) that would be expected for
igneous rocks differentiated directly from the mantle, or
remelted or equilibrated at high temperatures with such
rocks (5.3±0.3&, Valley et al. 1998). A distinct subset,
described below, has higher values up to 7.4&.
Jack Hills
The Jack Hills are in the Narryer gneiss terrane, NW
Yilgarn craton, Western Australia. Conglomerates and
quartzites contain detrital zircons and have experienced
upper greenschist to amphibolite facies metamorphism.
The values reported in Appendix 1are the average d
18
O
for each of 57 detrital zircons from four samples of the
Jack Hills metaconglomerate and quartzite, including
sample W-74 that was previously known to contain
zircons with ages from 4.4 to <3.1 Ga (Compston and
Pidgeon 1986; Wilde et al. 2001). These zircons are in-
terpreted as igneous based on fine concentric or sector
zoning imaged by cathodoluminescence. Some zircons
contain inclusions of quartz and devitrified melt. The
ages that are less than 3.73 Ga match the known crys-
tallization history of granitic rocks and gneisses that
outcrop adjacent to the Jack Hills. However, no rocks
are known that are old enough to be the source of zir-
cons older than 4 Ga. Several zircons have now been
found older than 4.3 Ga, which are the oldest known
terrestrial samples (Wilde et al. 2001; Cavosie et al.
2004; unpublished data). The d
18
O values in Appendix 1
represent averages of up to 12 independent spot analyses
on a single zircon (Peck et al. 2001; Cavosie et al. 2005).
The analytical precision on a single ion microprobe spot
(20 lm dia.) in 2004 is ±0.2–0.6&(1 SD). The un-
certainty of the mean improves to ±0.1–0.2&for ten
replicate analyses in a homogeneous crystal.
These in situ ion microprobe analyses provide unique
information on intra- and inter-crystal variation that is
essential for correlation of d
18
O and age. However, ion
microprobe analysis of d
18
O in detrital zircons is a dif-
ficult and relatively new technique. The only tabulated
data published for critical evaluation come from two
studies that show CL images and describe a careful
protocol for analysis that is required for accurate stan-
dardization (Peck et al. 2001; Cavosie et al. 2005). As a
further test of accuracy, we compare the average d
18
Oof
all Jack Hills zircons in Appendix 1(d
18
O(Zrc)=
6.2±0.7&, range = 5.0–7.4&) to the d
18
O measured by
more accurate and precise laser fluorination for a bulk
sample of several hundred Jack Hills zircons
(6.3±0.1&, Peck et al. 2001). These average values are
identical within analytical uncertainty to each other and
the range for Jack Hills zircons is the same as for all
other Archean igneous zircons that have been analyzed.
Higher values of d
18
O from 7 to 15&have been re-
ported for ion microprobe analyses of at least 15 pre-
4 Ga zircons from the Jack Hills (Mojzsis et al. 2001;
Trail et al. 2005), but neither the data tables nor images
necessary for interpretation of this unusual result are
published yet. These analyses are not included in Fig. 4.
If magmatic, values above 7.5&would be in contrast to
the Archean results summarized here (see discussion in
Peck et al. 2001; Cavosie et al. 2005). Cavosie et al.
(2005) report one 3.9 Ga zircon with d
18
O = 10.3, which
is interpreted as altered based on crosscutting textures
seen in CL (their Fig. 7). Because of the complexity seen
in CL images of some zircons and the difficulty of stable
isotope analysis by ion microprobe, such results require
detailed documentation.
Beartooth Mountains
Ten zircons were analyzed for d
18
O by ion microprobe
from the Beartooth Mountains, Montana in the
50
40
30
20
10
0
0 1 2 3 4 5 6 7 8 9 10 11 12 1413
δ18O (Zircon) ‰ VSMOW
Frequency
Average δ18O = 6.35 ± 1.56‰
N = 637
Phanerozoic
30
25
20
15
10
5
0
Frequency
Average δ18O=7.26±1.55‰
N=366
Proterozoic
30
25
20
15
10
5
0
Frequency
Average δ18O = 5.82 ± 0.74‰
N = 190
Archean
a
b
c
Fig. 3 Histograms of d
18
O(Zrc) for igneous zircons: aArchean; b
Proterozoic; cPhanerozoic
566
Archean Wyoming Province. These detrital zircons are
from quartzites in the same area of Hell Roaring Plateau
where Mueller et al. (1992,1998)reportedSHRIMPages
as old as 4.0 Ga. Our new SHRIMP data confirm this age
distribution and nine zircons between 4.0 and 3.7 Ga were
analyzed for d
18
O. The values of d
18
O average 6.2±0.5&
and range from 5.5 to 7.3&, nearly identical to the older
detrital zircons from the Jack Hills.
Barberton
Zircons were analyzed from 11 rocks from Barberton,
South Africa (3.5–2.7 Ga). The samples average
d
18
O(Zrc) = 5.5±0.7&(King 2001). These are the
oldest zircons that have been analyzed by high precision
laser fluorination. The values are indistinguishable from
the younger Superior Province samples.
Superior Province
The largest set of Archean d
18
O(Zrc) data comes from
58 plutonic rocks and 44 volcanic rocks of the Superior
Province of Canada (3.0–2.7 Ga). These samples are
representative of over 100,000 km
2
within the southern
SiO2 (Wt. %)
14
12
10
8
6
4
2
5040 60 70 80
δ18O (Zircon) ‰ VSMOW
Superior Province
3.0 to 2.7 Ga
Fig. 5 Plot of d
18
O(Zrc) versus SiO
2
content for magmatic zircons
and their host rocks for 35 samples from the Archean Superior
Province, Canada. Most samples are tonalites, trondhjemites and
granodiorites (TTG) or associated volcanic rocks and average
d
18
O=5.5±0.7&(1 SD). High-Mg sanukitoids have mild enrich-
ments in
18
O/
16
O, averaging 6.5±0.4&
Fig. 4 Compilation of d
18
O(Zrc) versus age for zircons from 1,200
rocks with known age. Samples range in age from 4.4 Ga to 0.2 Ma
and come from many terranes on seven continents. A remarkable
uniformity is seen in the Archean, values cluster near the mantle
(d
18
O(Zrc) = 5.3±0.3&) with some values as high as 7.5&
(horizontal line) due to recycling of supracrustal material into
magmas. Higher d
18
O, above 7.5&, only occurs after 2.5 Ga,
reflecting intra-crustal recycling of high d
18
O material and
maturation of the crust. Oxygen isotope data are from Table 1
and Appendix 1. Periods of supercontinent growth are shown by
short bars at bottom:PPangea; GGondwana; RRodinia, EP Early
Proterozoic; AArchean; and MA Middle Archean (Condie 1998).
LTB Late Terminal bombardment, CEE Cool Early Earth (Valley
et al. 2002), and O
2
= rise of oxygen in the atmosphere
567
and western Superior Province (Fig. 6). Of the plutonic
rocks, 42 are pretectonic to syntectonic, mostly tonalite-
trondhjemite-granodiorite (TTG) and d
18
O averages
5.6±0.5&(1SD, Fig. 7, King et al. 1998; King 2001).
Volcanic zircons yield the same values, d
18
O(Zrc)=
5.4±0.8&(King et al. 1997;2000; King 1997). In con-
trast, higher d
18
O values of 6.5±0.4&(n=17) come
from late tectonic to post-tectonic (2.70–2.68 Ga), Mg-
and LREE-enriched plutons with sanukitoid affinities
(King et al. 1998). These mildly elevated values support
the model that sanukitoids formed during subduction by
melting of altered upper ocean crust and/or peridotite in
the overlying mantle wedge that was metasomatized by
fluids from ocean crust (Shirey and Hanson 1984; Stern
and Hanson 1991). The upper portion of ocean crust is
elevated in
18
O/
16
O due to low temperature hydro-
thermal alteration and the presence of high d
18
O sedi-
ments, and thus metasomatizing fluids are likewise high
in d
18
O (Eiler et al. 1998; Eiler 2001). Some TTG plu-
tons from metasedimentary belts also show mildly
higher d
18
O. For instance, three plutons from the me-
tasediment-rich English River subprovince have
d
18
O(Zrc)=6.6±0.2&recording crustal contamination
(King et al. 1998).
Other rocks of the same age, 3.0–2.7 Ga, have similar
values including: two samples of tonalitic gneiss from
the Lewisian of Scotland (d
18
O(Zrc) = 5.5), five samples
from the Slave province, Canada (d
18
O = 4.9±0.3&,
King et al. 2001), and two samples from the Ukraine
that were analyzed in conventional nickel reaction ves-
sels (Lugovaya et al. 2001).
There are three Archean low d
18
O outliers in Fig. 4
at 2.7 Ga that are from volcanic rocks. Two samples
were collected from cordierite-orthoamphibole meta-
volcanics in the stockwork zone of the Manitouage
volcanogenic massive sulfide deposits at Geco, Ontario.
In this environment, low d
18
O values are common due
to the high temperature alteration of basalt by sea-
water, which formed the Mg-Al-rich protoliths to cor-
dierite gneisses (Pan and Fleet 1995). The low d
18
O
values of igneous zircons (2.4 and 4.4&) show that
these low d
18
O protoliths were remelted (Peck 2000).
The third low d
18
O zircon (2.9&) is from rhyolite at
Winston Lake and is explained by remelting of altered
rocks in the Sturgeon Lake caldera complex (King
et al. 2000).
Proterozoic
The d
18
O for zircons from a total of 366 Proterozoic
rocks range from 1 to 13&(average 7.3±1.5&). This
average is 1.4&higher than for all of the Archean
samples and the range is two to three times larger
(Fig. 3b vs. 3a). The values are bimodal. Similar differ-
ences (Fig. 7) were reported earlier for zircons from the
Grenville Province vs. the Superior Province (Peck et al.
2000). It is now clear that this evolution from uniform
d
18
O(Zrc) values (5 to 7.5&) in the Archean to more
variable and higher values (5to>10&) in the Pro-
terozoic occurs worldwide and is not restricted to N.
America.
Grenville Province
The most heavily studied and most variable Proterozoic
samples are from the Grenville Province (ca. 1.35–
1.0 Ga: Adirondacks, NY; Ontario including the Fron-
tenac terrane; and Quebec; Valley et al. 1994; Peck et al.
2004), and Grenville outliers in Vermont, the Blue
Ridge, Goochland regions of Virginia, and the Balti-
more Gneiss of Maryland. The zircons from 107 rocks
vary from 4.6 to 13.5&and average 7.9±1.6&(one
sample from a pegmatite intruding low d
18
O skarn is
1.0&).
Figure 8shows the ages and d
18
O values for the
Grenville Province. A major magmatic event, the ca.
Fig. 6 Map of the Superior and Grenville Provinces of North
America showing sample localities (after Peck et al. 2000)
Fig. 7 Histograms of d
18
O for igneous zircons from the Superior
and Grenville Provinces (after Peck et al. 2000)
568
1.15 Ga AMCG suite (anorthosite–mangerite–char-
nockite–granite), is seen for plutons from the Adir-
ondacks and southern Grenville. Figure 7shows d
18
O
for zircons from Grenville plutons: pre-AMCG (1.34–
1.18 Ga), AMCG (1.18–1.13 Ga), and post-AMCG
(1.09–1.05 Ga) (Peck et al. 2000,2004).
Figure 9shows the Grenville-age d
18
O(Zrc) values
versus whole rock SiO
2
. The majority of samples fall
between 6 and 10&and d
18
O shows no correlation with
SiO
2
, which varies from 41 to 77 wt.%. This range of
d
18
O values is representative of the entire southern
Grenville province and is higher than seen in Archean
samples.
A group of eight samples have unusually high
d
18
O(Zrc) of 11 to 13&, corresponding to the spike at
1.15 Ga in Figs. 4and 8. These samples are from a
relatively small group of AMCG-age plutons in the
Frontenac arch and NW Adirondack Lowlands between
Ontario and the central Adirondack Highlands, NY.
Silica varies from 61 to 75 wt % for these rocks (Fig. 9).
The Frontenac granites were first identified as high in
d
18
O by Shieh (1985) from whole rock data and have
been intensely studied because of their unusual oxygen
isotope ratio (Marcantonio et al. 1990; Peck et al. 2004).
The new zircon analyses show that these high d
18
O zir-
cons crystallized from high d
18
O magmas and are not the
result of postmagmatic alteration. These are the highest
d
18
O igneous zircons that have been reported and their
compositions are anomalous in Figs. 3b, 4,7,8, and 9.
Such magmas must result from melting of sediments
and/or altered ocean crust, which were buried deeply,
probably during continent–continent collision at ca.
1.2 Ga (Peck et al. 2004). The unusually high values
(d
18
O(Zrc) > 10&) are only seen in the post-Elzevirian
AMCG suite. Regardless of their genesis, the number of
analyses of these rocks over-represents their volume in
the crust. The majority of Grenville crust is represented
by values of d
18
O(Zrc) from 6.0 to 9.5&. The values for
Superior Province zircons (Fig. 5) are outlined in Fig. 9
emphasizing the contrast in d
18
O between these adjacent
terranes.
The Grenville-age samples (1.1–1.0 Ga) from Virgi-
nia (Blue Ridge, Goochland) and Maryland (Baltimore
Gneiss) are less variable in d
18
O than the samples from
the Adirondacks of New York and adjacent terranes in
Ontario but still are significantly higher (average 7.4&)
than in the Archean.
Finland
Oxygen isotope ratios of igneous zircons from granitoids
that intrude the Svecofennian of Finland also reveal
discontinuities in the deep crust. Three magmatic source
regions with distinct oxygen and neodymium isotope
signatures are revealed in a north–south traverse. Zircons
from the 1.88–1.87 Ga Central Finland Granitoid
Complex (CFGC) range from 5.5 to 6.8&(n=7), except
for three plutons in contact with supracrustal belts.
South of the CFGC, zircon from 1.65 to 1.54 Ga rapa-
kivi granites average 8.1±0.6&(n=5). Lastly, zircons
from 1.65 to 1.54 Ga granites in southernmost Finland
average 6.1±0.1&(n=3). These three magmatic source
regions are interpreted to reflect differences in accreted
Paleoproterozoic island arc terrains (Elliott et al. 2005).
South America
Zircons from Proterozoic rocks in South America were
analyzed from Brazil and Uruguay. Ages fall into two
groups: 2.36–1.70 Ga and 653–560 Ma. Early Proter-
ozoic samples are from the Ribeira belt, the Curitiba
microplate, Luis Alves microplate, SE border of San
Francisco craton, Caldas Branda
˜o massif, and the Pie-
dra Alta terrane. Late Proterozoic samples are from NE
Parana State, Paranagua
´and Monguagua
´batholiths,
Serra do Mar Alkaline-Peralkaline suite, Pien batholith,
Floriano
´polis batholith, Pelotas batholith, Sao Rafael
pluton, Serido terrane, Emas pluton, Cachoeirinha ter-
rane, Aigua batholith, and the Lavalleja metamorphic
complex. Data for these samples are reported in Table 1
SiO2 (Wt. %)
14
12
10
8
6
4
2
5040 60 70 80
δ18O (Zircon) ‰ VSMOW
Grenville - age, N. America
Superior Province
Frontenac
1.3 to 1.0 Ga
Fig. 9 Plot of d
18
O for magmatic zircons versus SiO
2
content of
their host rocks for 75 samples from the Grenville Province.
Frontenac samples are shown as open boxes. The field for Superior
Province samples is shown for comparison from Fig. 5
14
12
10
8
6
4
21000 1100 1200 1300
Age (Ma)
δ18O (Zircon) ‰ VSMOW
Grenville - Age, N. America
AMCG
Fig. 8 Plot of d
18
O(Zrc) versus age for magmatic zircons from the
Grenville Province. High d
18
O rocks from the Frontenac Arch with
d
18
O>10&formed during the AMCG magmatism at 1,180–
1,130 Ma (vertical dashed lines)
569
and Appendix 1(Ferreira et al. 2003). Appendix 2shows
d
18
O(Zrc) versus wt.% SiO
2
for the Late Proterozoic
rocks from Brazil. As for other suites, there is no cor-
relation of d
18
O and SiO
2
.
Neoproterozoic
The Late Proterozoic plutons that intrude the Gren-
ville-age rocks in the Blue Ridge of Virginia have a
measurably lower average d
18
O than nearby Grenville-
age plutons (average 6.4 vs. 7.4&) and suggest addi-
tion of juvenile magmas within an evolved high d
18
O
province.
Other Terranes
Zircons from smaller suites of Proterozoic samples were
analyzed from: northern Australia; basement in the
Basin and Range, western US.; Trans-Hudson, Canada;
Nain Anorthosite Complex, Canada; Laramie anortho-
site complex, Wyoming (O’Connor and Morrison 1999);
Ukraine (Lugovaya et al. 2001); and Penokean of Wis-
consin. The ages, values of d
18
O, and references for these
samples are summarized in Table 1and tabulated in
Appendix 1.
Phanerozoic
Sierra Nevada
The Sierra Nevada batholith, USA, is dominated by
Cretaceous plutons intruded into predominantly Jur-
assic and Triassic granitoids, metasediments, and me-
tavolcanics. Zircons have been analyzed for d
18
O from
287 rocks varying in age from 143 to 74 Ma, and 40
rocks from 222 to 145 Ma (Lackey 2005; Lackey et al.
2005a,b). Values of d
18
O(Zrc) are highly variable with
no significant difference between Cretaceous and Jur-
assic/Triassic plutons (7.0±0.9&and 6.7±0.7&, re-
spectively).
Consistent differences in d
18
O are seen correlating to
location within the batholith, rock type, and depth of
emplacement. The highest d
18
O zircons are from 35
samples from peraluminous garnet-bearing plutons,
which average 7.9±0.5&(Lackey et al. 2005b). If the
peraluminous rocks are not included, the difference
between Cretaceous and older granitoids is not sig-
nificant (6.8±0.8&vs. 6.7±0.7&). However, distinct
geographic differences persist between the eastern,
southern, and northern Sierra, and other areas (t-test at
greater than 99% confidence level): western Sierra,
6.8±0.8&n=36; central Sierra, 6.9 ±0.6&n=91;
eastern Sierra, 6.2±0.5&n=85; Owens Valley/ White
Mountains, 6.8±0.7 n=31; and northern Sierra,
5.5±0.6&n=3 (Table 1). While most of the Sierra
Nevada batholith presently exposes rocks that
intruded at depths of 5–13 km, the southernmost Sierra
(Tehachapi Mountains) intruded at 20–30 km and re-
present deeper levels of the batholith. Zircons from the
southern Sierra are the highest from metaluminous
gabbro, diorite, and tonalite plutons, and average
7.8±0.7&(n= 45), reflecting melting of metasedi-
mentary rocks.
It is intriguing in the Sierra Nevada that d
18
O(Zrc)
and initial
87
Sr/
86
Sr values have a negative correlation
over much of the batholith, and that lower values of
d
18
O(Zrc) are found in the east, towards the craton. In
fact, some of the highest d
18
O(Zrc) values are from rocks
with Sr
i
less than 0.706. The opposite relation is pre-
dicted for a west to east transition of oceanic to con-
tinental crust or for AFC processes involving high d
18
O
sediments. The adjacent Peninsula Ranges batholith
shows the predicted trends (Taylor 1986), in distinct
contrast to the Sierras. The negative correlation of d
18
O
and Sr
i
is evidence in the Sierras for considerable re-
cycling of young (Paleozoic or Mesozoic), hydro-
thermally altered upper oceanic crust or volcanic arc
sediments within the arc (Lackey et al. 2005a; Lackey
2005). The occurrence of lower average d
18
O in grani-
toids of the eastern Sierra, on the cratonic side of the
arc, indicates that magmas there were derived from aged
lithospheric mantle and were not significantly con-
taminated by overlying craton-derived sediments
(Lackey 2005).
Great Basin, Western US
Zircons have been analyzed from 124 Jurassic to Ter-
tiary granitic rocks from the Great Basin of Nevada and
Utah (King et al. 2004). Samples span mapped isopleths
for
87
Sr/
86
Sr
i
= 0.708 and 0.706, and Nd = 7.
Zircons of all ages show an increase in d
18
O to the east
of the 0.706 line, correlating with increased ratios of
whole rock Al
2
O
3
/(CaO + Na
2
O+K
2
O). The crustal
boundaries defined by radiogenic isotopes in the Great
Basin agree with discontinuities in d
18
O(Zrc) in contrast
to whole rock d
18
O values, which are frequently altered
and correlate poorly.
Idaho Batholith
The late Cretaceous and Tertiary granitic rocks of the
Idaho batholith intruded the Precambrian margin of
North America. Values of d
18
O(Zrc) are relatively
homogeneous in spite of prolonged magmatic history.
Zircons in the Bitterroot Lobe (northern part of bath-
olith) average 7.1±0.3&(n= 7), while in the Atlanta
Lobe (southern), they average 6.7±1.5&(n=19). Eo-
cene plutons average 7.2±0.2&(n= 3) with one ex-
ception at 4.0&(King and Valley 2001).
Eastern China
A-type granites from four late-Mesozoic plutons in
eastern China have an average d
18
O(Zrc) = 4.9±0.3&
570
(n=30), while a fifth pluton averages d
18
O(Zrc) = 3.7
±0.4&(n= 6) (Wei et al. 2002; unpublished). These
mildly low d
18
O magmas suggest protoliths or magmatic
contaminants that exchanged with surface waters at high
temperature.
British Tertiary Igneous Province
Sub-volcanic igneous centers have been studied from the
Isles of Skye, Arran, and Mull in Scotland demon-
strating the presence of low d
18
O values as a result of
magmatic and post-magmatic processes typically loca-
lized within eroded caldera complexes. In spite of ex-
treme hydrothermal alteration of many rocks, all
evidence indicates that low d
18
O values in zircon are
magmatic compositions. In many cases, the low d
18
O
magmas resulted from cannibalization, i.e., remelting of
hydrothermally altered earlier phases of the same ig-
neous suite (see, Valley 2003). The resulting d
18
O(Zrc)
values range from 0.6 to 7.1&(n=27, Gilliam and
Valley 1997; Monani and Valley 2001).
Tertiary volcanic rocks, Western US
Volcanic rocks have been studied in detail from cal-
dera complexes at Yellowstone, Long Valley, and
Timber Mountain/Oasis Valley in the western United
States. Relatively small volume, postcaldera rhyolites
at Yellowstone have low d
18
O(Zrc) values to 0.0&,in
comparison to the large (600–2,500 km
3
) caldera
forming Huckleberry Ridge and Lava Creek tuffs
(d
18
O(Zrc) = 4.1–5.7&, Bindeman and Valley 2000,
2001). Low d
18
O rhyolites are also found at the
Timber Mountain/ Oasis Valley Caldera complex in
Nevada where smaller depletions of 1–2&are seen,
but the volumes of low d
18
O rock are significantly
larger for the Tiva Canyon and Ammonia Tanks tuffs
(900–1,000 km
3
, Bindeman and Valley 2003). In con-
trast to these nested caldera complexes, zircons from
the Bishop tuff (>650 km
3
) at Long Valley caldera
are mantle-like and homogeneous in d
18
O (5.8
±0.2&,n=4, Bindeman and Valley 2002).
Mantle zircons
Zircon megacrysts are a trace constituent in many
kimberlites. Typically, the U-Pb age matches the
eruption age of the kimberlite pipe, and the d
18
Oof
zircons approximates high temperature equilibration
with the mantle (d
18
O(Zrc)=5.3±0.3&). This value
would be in magmatic equilibrium with an oceanic
basalt at d
18
O(WR) = 5.5&and is the predicted va-
lue of d
18
O in primitive mantle-derived magmas. While
d
18
O is very homogeneous for zircons from each pipe,
within less than ±0.2&, small regional variability is
observed with some pipes having values either above
or below the mantle value (Valley et al. 1998; un-
published data). Ion microprobe analysis of a few
selected crystals, including KIM-4 and KIM-5 stan-
dards, has shown intra-crystalline homogeneity (Peck
et al. 2001; Valley 2003; Cavosie et al. 2005). Pre-
cambrian zircons from Zwaneng, Botswana are
anomalous and show inter- and intra-crystalline
variability (Valley et al. 1998; Valley and McKeegan
unpublished), consistent with a prolonged history in
the crust. The kimberlite zircons are a distinct suite
with clear mantle affinities and will not be considered
further in this paper, which addresses the maturation
of continental crust.
A Secular Change in magmatic d
18
O
Figure 4shows a secular change in d
18
O(Zrc). Zircons
from younger magmas are more variable and many are
higher in d
18
O. The large amount of information in
Fig. 4complicates the simple trend and has been re-
plotted in Fig. 10a where all data have the same
symbol. This figure emphasizes that values were rela-
tively low and constant throughout the Archean, and
shows that the trend towards increasing values began
at 2.5 Ga. A horizontal line at d
18
O = 7.5&defines
the highest values in the Archean. After 2.5 Ga, the
upper limit of d
18
O(Zrc) increases to 10&at ca. 1 Ga,
encompassing all data except the anomalous Frontenac
samples at 1.15 Ga. While it might be tempting to fit a
more complex curve to these data with peaks and
valleys, the valleys fall in intervals with less data and
probably result from the statistics of small popula-
tions.
The interpretation of the trend in Fig. 4depends
critically on the conclusion that all values are faithfully
preserved from the original magma. Apparent secular
trends of d
18
O in carbonates and cherts have been at-
tributed by some to problems of preservation, where
the oldest samples are interpreted to be most altered.
There are two reasons why the trend in Figs. 4and 10a
cannot be dismissed as the result of postmagmatic
disturbance. First, all evidence indicates that crystalline
zircons reliably preserve their magmatic value of d
18
O,
and second, there is no reason why alteration would
affect only the youngest rocks. Disturbance of d
18
O
would be expected to be greater on average in older
rocks, which have had more opportunity to experience
metamorphism, radiation damage, and other forms of
alteration. Clearly, the trend in Figs. 4and 10a is the
opposite of that expected if older samples are more
altered.
It is important to emphasize that there are no known
primitive reservoirs in the mantle for the extreme
d
18
O(Zrc) values of Fig. 4(e.g., >6.0&or <4.6&).
571
Fractionations are small at high temperatures and thus
formation of d
18
O values higher or lower than the
mantle requires protoliths that were altered near the
surface of the Earth where temperatures are low and
fractionations are large. Where anomalous d
18
O values
have been identified in the mantle, they represent sub-
ducted supracrustal material.
Thus, the variability of d
18
O(Zrc) is a sensitive re-
cord of recycling of supracrustal lithologies. Figure 1
shows that reasonable supracrustal contaminants range
in d
18
O to values above 25&. Even small amounts of
such near-surface rocks would cause a measurable in-
crease in d
18
O of magmas and zircons. For instance,
5% contamination by a rock with d
18
O=25&would
raise d
18
O in a normal magma by 1&. Likewise, bulk
melting of igneous rocks mixed with 5% metasediments
could raise d
18
Oby1&. Thus, while questions remain
about timing and cause, it is apparent that processes of
intracrustal recycling have systematically changed the
amounts of contamination and melting through time
and/or the d
18
O of fertile crust has increased through
time.
High d
18
O magmas
Values of d
18
O(Zrc) above 7.5&are common in mag-
mas younger than 1.5 Ga. These high d
18
O zircons are
representative of rocks outcropping over large areas of
several terranes, including: the southern Grenville Pro-
vince (1.34–1.05 Ga, Fig. 7); Grenville-age outliers in
the Appalachians; several regions of South America
(Emas and Tavares plutons, 650 Ma); and the Sierra
Nevada batholith (222–74 Ma). Smaller proportions of
other terranes have high d
18
O igneous compositions
including: the Ukrainian shield (2.62–1.93 Ga); Aus-
tralia (1.86–1.80 Ga); Finland (1.88–1.57 Ga); the Great
Basin (1.5 Ga); and the Idaho batholith (94–70 Ma).
With the exception of the Frontenac AMCG suite and
garnet-bearing peraluminous granitoids of the Sierra
Nevada batholith, these high d
18
O rocks were not tar-
geted for special study. Thus, the proportion of high
d
18
O zircons in Appendix 1is the best estimate available
for the proportion of high d
18
O igneous rocks in these
areas.
The high d
18
O(Zrc) values are in contrast to the
Archean when values are lower, close to the d
18
O(Zrc)
values expected in mantle-derived (4.6–6&) or mildly
evolved (6–7.5&) magmas. Zircons from 7.5–10&in-
dicate high d
18
O whole rock values of 9 to 12&for
Fig. 10 Mantle evolution and crustal growth, and maturation are
shown by stable and radiogenic isotopes. a. Values of d
18
O(Zrc)
and age of igneous zircons (Fig. 4). b. Distribution of zircon ages
from juvenile crust (Condie 1998). 10c. Crustal growth curves
representative of many that have been proposed based on the
preserved rock record and a simplified line from Kramers (2002). d
and e. Evolution of the depleted mantle as shown by Nd and Hf
isotopes compiled by Bennett (2003) who infers a rapid period of
early fractionation followed by steady state through much of the
Archean and major fractionation after 2.5 Ga
b
572
felsic magmas, consistent with derivation by melting
of sediments as is observed in ‘‘S-type’’ granites
(Fig. 1, O’Neil and Chappell 1977; O’Neil et al. 1977;
Taylor and Sheppard 1986). However, many of the
host rocks for these high d
18
O zircons are not per-
aluminous and major element chemistry is not a good
predictor for high d
18
O magmas in the absence of
other evidence. Low temperature alteration is a com-
mon mechanism that can raise the d
18
O of near-sur-
face rocks without necessarily changing other chemical
characteristics. Thus, it is likely that assimilants other
than high-Al clays are important and that oxygen
isotope behavior is variably decoupled from other
geochemical systems.
The number of high d
18
O(Zrc) samples increases
gradually over a period of approximately one billion
years during the Proterozoic (Figs. 4,10a). The one
exception to a smooth trend is the anomalous group of
high d
18
O Frontenac samples. The transition from lower
to higher d
18
O starts at approximately the end of the
Archean, but the exact timing is poorly constrained due
to limited samples between 2.7 and 2.0 Ga.
This secular change in oxygen isotopic reservoir
characteristics marks a major non-uniformitarian tran-
sition in the Earth’s continental crust. Higher d
18
O va-
lues result from subduction or burial of high-d
18
O
sediments, and rocks weathered or altered at low tem-
peratures, which are then recycled in high d
18
O magmas.
Clearly, such high d
18
O rocks have been increasingly
recycled within the crust starting in the Proterozoic.
Low d
18
O magmas
Young samples with d
18
O<5&are common in Fig. 4
for the past 150 Ma, but only a few scattered samples
are seen from older rocks. A few values in this time
period are as low as 0&. These data are tabulated in
Appendix 1, but values below 2.3&are not shown in
Fig. 4to save space. It is reasonable to ask if the con-
centration of values that are lower than the mantle also
represent a secular change in magmatic d
18
O. The low
d
18
O samples represent three relatively small areas that
were chosen for close examination specifically because
previous studies suggested the presence of low d
18
O
magmas: sub-volcanic granites from the British Tertiary
Igneous Province (Gilliam and Valley 1997; Monani and
Valley 2001) and eastern China (Wei et al. 2002; un-
published); and low-d
18
O rhyolites from Yellowstone
(Bindeman and Valley 2001,2002). These three areas
represent shallow sub-volcanic magma chambers where
low d
18
O values resulted from melting of hydrothermally
altered wall rock. In the course of investigations of the
genesis of low d
18
O magmas, the lowest d
18
O rocks from
each area were intensely collected and studied, biasing
the proportions of analyses in Figs. 3c and 4. For in-
stance, at Yellowstone, the largest caldera-forming
eruption, Huckleberry Ridge tuff was 2,500 km
3
in vo-
lume and the low d
18
O rhyolites represent intra-caldera
flows of 10–50 km
3
each (Hildreth et al. 1984), yet the
number of analyses for low d
18
O rhyolites exceeds that
for the caldera-forming eruptions. Thus, the volume of
low d
18
O magmas is significantly over represented by the
number of analyses in Figs. 3c and 4.
Extremely low d
18
O zircons (to 11&) of igneous
origin are reported from the Dabie-Sulu terrane, China
(Rumble et al. 2002; Chen et al. 2003; Zheng et al. 2004;
Zhao et al. 2004). These zircons have been intensely stu-
died because host rocks include coesite- and diamond-
bearing eclogites. Igneous ages are 0.8 to 0.7 Ga, 0.5 Ga
older than ultrahigh pressure metamorphism. The low
d
18
O values are attributed to high temperature exchange
with very low d
18
O glacial melt-water during Snowball
Earth events in the Neoproterozoic (Rumble et al. 2002;
Zheng et al. 2004). No other zircons below 0&have
been reported and these unusual values may be unique.
They are not plotted in Fig. 4for simplification.
The question persists; why do so few magmas have
d
18
O(WR) below 6&before 150 Ma? One could answer
that examples are more common than is recognized and
could be located by directed study. For instance, the low
d
18
O samples from the Superior Province were found by
targeting the stockwork feeder zone of a volcanogenic
massive sulfide deposit and low d
18
O magmas are par-
ticularly abundant in volcanic areas undergoing glacia-
tion such as Kamchatka where glacial melt-waters have
very low d
18
O (Bindeman et al. 2004). In contrast,
Balsley and Gregory (1998) propose that the genesis of
low d
18
O magmas is a relatively rare event. While this
may be correct, it does not address the increasing rarity
for older terranes. Preservation must also be an im-
portant factor in the scarcity of old, low d
18
O igneous
rocks. Hydrothermal alteration by surface waters is re-
stricted to the relatively shallow crust, above the brittle-
ductile transition. Such near-surface rocks are pre-
ferentially eroded. Thus, low d
18
O magmas have always
been a feature on Earth, their volume has never been
great, and they have been selected against in the rock
record. There is no indication at present that the volume
of such rocks has either increased or decreased system-
atically through time.
Constant d
18
O in the Archean Crust
The uniformity of d
18
O values throughout the Archean
since 4.2 Ga is one of the most surprising and significant
findings of this study. For approximately two billion
years, the d
18
O of most igneous rocks averaged exactly
the mantle value, a smaller number of rocks have 1–2&
higher values, and no magmatic zircons are found with
d
18
O > 7.5&. The data-rich histogram for 2.7 Ga Su-
perior Province zircons (Fig. 7a) shows the same range
of d
18
O as the smaller sample set for Barberton. The less-
precise ion microprobe analyses of Early Archean zir-
cons from the Jack Hills and the Beartooths are slightly
higher on average, but within analytical uncertainty of
these values.
573
Significant differences exist even within the rela-
tively restricted and constant 2.5&range of Archean
d
18
O(Zrc). Values above 6&cannot be explained as
pristine differentiates from mantle magmas. To be
conservative, the lower limit of the non-mantle su-
pracrustal field is set at 6.5&when poorer precision
of ion microprobe data is discussed (vs. >6.0&for
laser fluorination data). These higher d
18
O values in-
dicate intracrustal recycling of surface materials into
magma by melting or contamination. The ultimate
source of higher values in the supracrustal materials
was from low temperature interaction with water in
the near-surface environment. Thus, the zircon record
indicates that igneous rocks in the crust achieved small
amounts of differentiation by 4.3 Ga and oxygen
isotope ratios maintained a steady state from 4.2 to
2.5 Ga.
Models for growth of the continental crust and rates
of recycling via subduction vary greatly (Fig. 10b–e, see,
Hurley and Rand 1969; Taylor and McLennan 1985;
Armstrong 1981;1991; Bowring and Housh 1995;
Condie 1998; Kramers 2002; Bennett 2003; Campbell
2003). In ocean crust, the intensity of hydrothermal al-
teration may have been greater in the Archean, but the
combined effects of high and low temperatures of ex-
change balanced each other such that no measurable
shift in average d
18
O occurred (Muehlenbachs 1998). In
continental crust, processes of crustal growth added
magmas with near-mantle d
18
O values, while intra-
crustal recycling of supracrustal rocks created magmas
with higher d
18
O (Simon and Lecuyer 2002).
It is remarkable that the steady state reflected by
d
18
O of magmas corresponds to the main periods of
crustal growth. This uniformity suggests that feedback
mechanisms operated throughout the Archean. The
creation of new continental crust is one consequence
of thermal events that are accompanied by heating
and remelting of existing crust (Kemp and Hawkes-
worth 2003). The oxygen isotope record shows that
throughout the Archean, rates of crustal growth were
balanced by rates of magmatic recycling in continental
crust. In contrast, the trend towards higher d
18
O va-
lues, which begins at the end of the Archean, results
from non-uniformitarian changes that altered the Ar-
chean steady state that had persisted for approxi-
mately two billion years. The rate of crustal growth
declined and the effects of intra-crustal recycling in-
creased.
Causes of secular change in the Proterozoic
The early Proterozoic was a time of great change on
Earth with increased sedimentary environments follow-
ing the period of crustal growth and cratonization in late
Archean. Several important transitions occurred affect-
ing: the composition and abundance of sedimentary and
igneous rocks available for recycling; the rates of sub-
duction; and differences in weathering as the atmosphere
became more oxygen-rich and life flourished. These
changes correlate in time to the major shift in d
18
O(Zrc)
and to trace element compositions (see Veizer and
Mackenzie 2003; Kemp and Hawkesworth 2003;
McLennan et al. 2005).
Sediments
Sediments are the dominant reservoir of high d
18
O
material on Earth (Fig. 1). The quantity and d
18
Oof
sediments available for burial and recycling impacts the
composition of resultant magmas. Thus, any process
that changes the d
18
O of sediments, or changes the
quantity of sediments available for melting, will have a
corresponding effect on the d
18
O of igneous rocks and
their zircons.
Figure 11 shows estimates for the evolving percentages
of different sedimentary rock types through time (Veizer
and Mackenzie 2003). Estimates are hypothetical before
3.0 Ga due to the incomplete rock record; later trends,
based on more data, support the conclusion that Archean
sediments were on average lower in d
18
O. Archean sedi-
ments are dominated by greenstone-belt sequences, which
are comprised largely of lower d
18
O volcaniclastic, pyr-
oclastic, and sedimentary material of non-cratonic origin
(Lowe 1994; Veizer and Mackenzie 2003). Continental
and continental margin sequences were not common be-
fore 3.5 Ga and the Early Proterozoic marked a transition
to major cratonic sedimentary sequences including an
increase in high d
18
O clays and chemical sediments.
Shales comprise the largest high d
18
O sedimentary
reservoir in the modern crust, but the fine-grained clastic
sediments that are observed in the Archean are less ma-
Limestones Jaspilites and
their analogues
Shales and metamorphic equivalents
Secondary
quartzites
Evaporites
Dolomites
Quartz
sands Arkoses
Graywackes
Submarine volcanogenics
Continental
extrusives
Volume %
100
75
50
25
00 300020001000 4000
A
g
e
(
Ma
)
Fig. 11 Volume percent of different sedimentary rock types as a
function of age. (from Veizer and Mackenzie 2003)
574
ture, richer in unaltered volcanic material, and lower in
d
18
O (Longstaff and Schwarcz 1977; Shieh and Schwarcz
1978; Peck et al. 2000). Increasing maturity and clay
content in clastic sediments have raised the bulk d
18
O
through time. Furthermore, Veizer and Jansen (1985)
present Sm-Nd model ages for sediments and conclude
that the quantity of sediment increased through the Ar-
chean by erosion of relatively young igneous rocks within
250 million years of differentiation from the mantle.
These ‘‘first cycle’’ rocks built up to nearly the present
mass of sediment by 2.5 Ga and recycling of sediments
then became dominant (Windley 1995). Thus, while
shales became quantitatively important after 3.5 Ga, the
proportion of second generation shales with higher d
18
O
increased after 2.5 Ga. Furthermore, in the Archean,
aggressive weathering in CO
2
-rich atmospheres may have
stripped sediments of feldspar, leaving quartz-rich clastic
rocks lacking in components that commonly make high
d
18
O clays (Lowe and Tice 2004). Other changes in
weathering pattern resulted from rising atmospheric
oxygen levels at ca. 2.3 Ga that created more oxidized
sediments and biological colonization of land. These
changes facilitated weathering of primary feldspars and
volcanic glass to form clays, which can be up to 30&
higher in d
18
O than coexisting surface waters (Savin and
Epstein 1970). Increased sedimentary reworking also
contributed to increased marine deposits that were sub-
sequently subducted. These long-term trends in weath-
ering and clastic sedimentation contributed to the
increase of magmatic d
18
O values during the Proterozoic.
Thus, the early Proterozoic saw increased amounts of
higher d
18
O sediments due to sedimentary recycling, and
the growth of continents and epicontinental seas (Veizer
1983; Taylor and McLennan 1985; Eriksson 1995;
Windley 1995; Condie et al. 2001).
Secular increases of 5–10&in the d
18
O of chemical
sediments, limestones and cherts, are also reported during
the Proterozoic that would further contribute to the
change in magmas (Shields and Veizer 2002; Perry and
Lefticariu 2003; Knauth and Lowe 2003; but see Land
and Lynch 1996 for shales). The causes and significance of
these trends are controversial. Competing interpretations
propose a secular increase in d
18
O of the younger oceans;
higher ocean temperatures in the Archean, or greater al-
teration of older sediments. The largest proposed changes
in d
18
O of the ocean (>10&) are not consistent with the
compositions of igneous rocks altered by seawater
(Muehlenbachs 1998), but smaller differences might not
be discerned. Likewise, the highest proposed ocean tem-
peratures (‡70C) must be reconciled with Precambrian
continental glaciations, but smaller, localized, or inter-
mittent increases in temperature could still lead to lower
d
18
O sediments in earlier rocks and higher values in the
Proterozoic. In addition to evolving d
18
O values, the
quantities of high d
18
O carbonates and other chemical
sediments are greater in the Proterozoic (Windley 1995;
Veizer and Mackenzie 2003). Any combination of these
processes would contribute to the secular change in
magmas as seen in the zircon record.
Burial ± subduction
Burial is prerequisite for supracrustal rocks, whether
sedimentary or volcanic, to be melted or assimilated by
magma. For at least the past 2.5 Ga, subduction has
been an important process to bury rocks and generate
silicic and intermediate composition magmas, and some
form of plate tectonics may have started much earlier (de
Wit 1998). Nevertheless, other processes were operative
in the Archean that do not require convergent tectonism
(Bleeker 2002). Thick volcanic successions in extensional
or plume-dominated environments can lead to burial
and melting of sediments or igneous rocks. In a modern
setting, caldera collapse and foundering of altered wall
rocks caused magmatic d
18
O to shift by several permil at
Yellowstone (Bindeman and Valley 2002). More ex-
tensive processes would have operated if Archean tec-
tonics were plume- and rift-dominated. The earliest
felsic crust may have formed on mafic basement in an
‘‘Iceland-like’’ environment (Kroner and Layer 1992)
with mildly elevated d
18
O values seen as early as 4.3 Ga
(Cavosie et al. 2005).
In contrast to plume tectonics, subduction reworks
greater amounts of crust both by melting of subducted
ocean crust with a sedimentary component from the
continents and by subsequent melting within the con-
tinental crust caused by metasomatism and magmatic
heating. Tectonic conditions were distinct in the Archean.
Higher radiogenic heat production fostered vigorous
greenstone-belt tectonics and many small unstable mi-
croplates. The average crust was younger and therefore
hotter at the time of subduction. Crustal material was
certainly returned to the mantle during the Archean, but
as radiogenic heat-production declined, the style of sub-
duction changed. In the Proterozoic, amounts of sub-
ducted sediment increased. Lowe (1992) has further
proposed that Himalayan-style subduction first occurred
in the late Proterozoic. If so, this process also could have
contributed to the dominant high d
18
O magmatism first
seen at 1.3 to 1.0 Ga in the Grenville Province. Thus, in-
creased rate and changing style of subduction are both
likely contributing causes of secular change of d
18
Oin
magmas.
The highest d
18
O values in Fig. 4show a steady in-
crease through the middle Proterozoic suggesting a
gradual build-up in the amounts of
18
O-enriched mag-
mas. Curiously, this trend has not been found to con-
tinue in the Phanerozoic. It may be significant that many
of the highest d
18
O Proterozoic rocks were metamor-
phosed at depths of 20–30 km. Either Earth reached a
new steady state with respect to oxygen isotopes at the
end of the Precambrian or, more likely, the younger high
d
18
O equivalents are not yet exhumed in great quantity
from deep in the crust. This later scenario is demon-
strated in the Sierra Nevada batholith; only a small
proportion of 4–9 kb granites are exposed and they are
systematically higher in d
18
O (Lackey et al. 2005a).
Alternatively, it is possible that d
18
O(Zrc) values above
10&are not common.
575
Thus, the secular rise of magmatic oxygen isotope
ratios through the Proterozoic is explained by a combi-
nation of changes in the composition, availability,
weathering, and burial of sediments that resulted from
tectonic changes at the end of the Archean. The details of
this important change will be elucidated as more geo-
chemical systems and techniques are employed to study
zircon and other retentive geochronology minerals.
Variation of d
18
O in the Mantle?
The dramatic trend to higher d
18
O in Proterozoic and
Phanerozoic magmas and the accumulation of high d
18
O
sediments and metasediments shows that the average
d
18
O of continental crust has increased from 4.4 Ga to
today. This increase must have been balanced by com-
pensating changes in mass or d
18
O of other reservoirs. To
identify these other reservoirs, a first order average d
18
O
of continental crust can be estimated by consideration of
sediments and magmas. Secondary processes of altera-
tion and metamorphism are also important, but a more
complete evaluation is beyond the scope of this paper and
is not required for this discussion. The d
18
O of sediments
range from 10 to over 40&and have been heavily studied.
Veizer and Mackenzie (2003) review the evolution of se-
dimentary rocks and summarize studies concluding that
14% of the continental crust is composed of sediments
with an average d
18
Oof17&. The d
18
O of magmas can be
estimated from Fig. 4taking account of the age dis-
tribution (Fig. 10c), the average fractionation between
zircon and whole rock (1&), and considering mafic
magmas which are lower in d
18
O (Harmon and Hoefs
1995) and not fully represented by zircon-bearing sam-
ples. As previously discussed, low d
18
O magmas are not
volumetrically significant. Taken together, the average
continental crust is estimated to be 9–10&. This re-
presents an elevation of d
18
O of ca. 4&relative to the
average mantle value of 5.5&.
A4&elevation in d
18
O for the entire continental crust
would require a major reservoir for mass-balance. Only
about 20% of this amount is compensated by low d
18
Oof
the oceans (0&). This leaves the mantle as the remaining
reservoir of sufficient size to compensate this change.
Mass-balance shows that if the entire 4&rise in
average d
18
O of the continental crust was compensated
by subduction of low d
18
O material into the upper
400 km of the mantle, the average decrease of d
18
Oin
the mantle over time would be approximately 0.1&. The
change in the mantle would be even less if subducted
material is mixed more deeply. The possibility that the
modern mantle is significantly heterogeneous in d
18
O,
due to failure of subducted material to mix on a suffi-
cient scale, has not been supported by analysis of peri-
dotites or oceanic basalts (Eiler 2001).
The best empirical evidence for d
18
O of zircon in
equilibrium with the primitive mantle reservoir in the
Archean comes from the Superior Province at 2.7 Ga.
The average value for zircons from TTG’s, volcanic
rocks, and other non-sanukitoid magmas is 5.5&.This
value is 0.2&above the value estimated for modern
magmas (Valley et al. 1998; Eiler 2001), suggesting ei-
ther that terrestrial granitoids are slightly evolved re-
lative to ocean basalts or that a secular change of 0.2&
has occurred over the past 2.7 billion years. While this
estimate is similar in magnitude and in the same direc-
tion as that predicted by mass-balance, the uncertainties
are relatively large. The earlier Jack Hills zircons appear
slightly heavier still, but better analytical precision by
new generation ion microprobes will be necessary to
evaluate this difference. Thus, while the secular change
in d
18
O of the crust is significant and most reasonably
balanced by subduction into the mantle, no detectable
change in the average d
18
O of the mantle is either pre-
dicted or resolved by the present analysis. A similar
conclusion was reached by Lowry et al. (2003) based on
analysis of olivine in 3.8 Ga ultramafic rocks. The mass
of the mantle is simply too large for its average d
18
Oto
be affected by the crust. More refined estimates are un-
likely to change this conclusion.
Supercontinent cycles
It has long been recognized that radiogenic isotope
ages are not uniformly distributed through time
(Fig. 10b) and much discussion has centered on the
question of whether magmatic differentiation and
growth of continental crust is a reversible process. One
view has held that once crust was created, it would
never be returned to the mantle, and the age spikes
(Fig. 10b) record the rate of crustal growth (Hurley
and Rand 1969). In contrast, Armstrong (see Arm-
strong 1981,1991; Sylvester 2000) argued for high rates
of sediment subduction, that the growth rate of crust is
underestimated by Hurley and Rand, and that growth
of the crust to present mass was effectively complete by
3.5 Ga (Fig. 10c). The Armstrong model thus posits
that there has been no net growth in the mass of
continental crust through time and that new additions
to the crust have been balanced by subduction in a
steady state. Both Nd and Hf isotopic data support
models of early differentiation of significant amounts of
continental crust (Figs. 10d+e), however these isotope
evolution diagrams do not resolve the question (Ben-
nett 2003). The oxygen isotope data for Archean zir-
cons provide a new constraint.
Recent models attribute the punctuated age dis-
tribution (Fig. 10b) to planet-scale cycles in the mantle,
related to supercontinent or superplume events (Stein
and Hoffman 1994; Condie 2000). Condie (2000) sum-
marizes arguments that episodic periods of super-
continent formation in the Precambrian and probably in
the Phanerozoic (Fig. 10) correlate to ‘‘catastrophic slab
avalanching at the 660 km discontinuity’’. Whatever
their causes, it is clear that these events have not altered
the steady state recorded by d
18
O of zircons in the Ar-
chean. It is likely that the additional heat advected into
576
the crust during magmatic pulses caused more melting
and recycling of the crust, but maintained a constant
proportion of primitive magma to remelted crust.
The changes in magmatic d
18
O starting at 2.5 Ga
support the models of Taylor and McLennan (1985) and
Kramers (2002) with high rates of growth for continental
crust in the Archean and significantly slower growth after
2.5–1.9 Ga. The oxygen isotope record is more difficult
to reconcile with the model of Armstrong (1981,1991),
which would produce higher d
18
O in Archean magmas,
or with that of Hurley and Rand (1969) that suggests
significant amounts of continental crust first appeared
after 1.5 Ga. If crustal growth was rapid from 4.4 to ca.
1.9 Ga, the ratio of primitive mantle magma to supra-
crustal material was relatively high, diluting amounts of
higher d
18
O magma and maintaining the subdued near-
mantle d
18
O values seen for Archean zircons. After the
major spikes of growth at 2.7 and 1.9 Ga, the propor-
tions changed. Significantly larger amounts of con-
tinental crust were available to be altered and to form
clastic and chemical sediments. At about the same time,
the rates of mantle magmatism declined. Over the fol-
lowing one billion years, intracrustal recycling of these
high d
18
O materials into magmas became increasingly
more important and created the secular trend seen in
Figs. 4and 10a. It is also possible that the absence of still
higher values (> ca. 10&) after 1 Ga represents estab-
lishment of a new steady state for oxygen isotopes in the
continental crust, reflecting the mass of continental crust
that was largely established by 1.9 Ga.
Acknowledgments We thank the following people who have pro-
vided samples, assisted, or collaborated in studies of these zircons:
John Aleinikoff, Tucker Barrie, Pat Bickford, Lance Black, Otto
van Breemen, James Carl, Jeff Chiarenzelli, Jim Chen, Fernando
Corfu, Louise Corriveau, Tony Davidson, Don Davis, John Eiler,
Brent Elliott, Ron Emslie, Dave Farber, Frank Florence, Carrie
Gilliam, Matthew Grant, Mike Hamilton, Hans Hinke, Martha
House, Yngvar Isachsen, Paul Karabinos, Yaron Katzir, Alan
Kennedy, Peter Kinny, Nami Kitchen, Bart Kowalis, Tom Krogh,
Dunyi Liu, Jim Mattinson, Jim McLelland, Dave Mogk, Salma
Monani, Sam Mukasa, Sasha Nemchin, Randy Parrish, Lola
Pereira, Bob Pidgeon, Helcio Prazeres Filho, Kent Ratajeski, Greg
Roselle, Jason Saleeby, Dan Schulze, Danny Stockli, Matti Vaas-
joki, Randy Van Schmus, Lee Silver, Sorena Sorensen, Beth Va-
laas, Julie Vry, Simon Wilde, Joe Wooden, and Jim Wright. Colin
Graham and John Craven collaborated in ion probe studies of d
18
O
at the Edinburgh Ion Microprobe Facility, which is supported by
NERC. Brian Hess aided with sample preparation. Mary Diman
drafted the figures. Vicki Bennett and Jan Kramers made helpful
reviews. This research was supported by the National Science
Foundation (EAR93-04372, 96-28142, 99-02973, 02-07340) and the
U.S. Department of Energy (93ER14389).
Appendix 1
Oxygen isotope ratio, crystallization age, and location
for magmatic zircons. Whole rock weight percentage
SiO
2
is tabulated where available. References to pre-
vious work include published and unpublished sources.
Table given as ESM, available at http://dx.doi.org/
10.1007/s00410-005-0025-8
Appendix 2
References
Armstrong RL (1981) Radiogenic isotopes: the case for crustal
recycling on a near-steady-state no-continental-growth earth.
Phil Trans Roy Soc Lond A 301:443
Armstrong RL (1991) The persistent myth of crustal growth.
Austral J Earth Sci 38:613–630
Balsley SD, Gregory RT (1998) Low-d
18
O silicic magmas: why are
they so rare? Earth Planet Sci Lett 162:123–136
Bennett VC (2003) Compositional evolution of the mantle. Treat
Geochem 2:493–519
Bindeman IN, Valley JW (2000) Formation of low-d
18
O rhyolites
after caldera collapse at Yellowstone, Wyoming, USA. Geology
28:719–722
Bindeman IN, Valley JW (2001) Low-d
18
O rhyolites from Yel-
lowstone: magmatic evolution based on analyses of zircons and
individual phenocrysts. J Petrol 42:1491–1517
Bindeman IN, Valley JW (2002) Oxygen isotope study of the Long
Valley magma system, California: isotope thermometry and
convection in large silicic magma bodies. Contrib Mineral
Petrol 144:185–205
Bindeman IN, Valley JW (2003) Rapid generation of both high-
and low-d
18
O, large-volume silicic magmas at the Timber
Mountain/Oasis Valley caldera complex, Nevada. Geol Soc Am
Bull 115:581–595
Bindeman IN, Ponomareva VV, Bailey JC, Valley JW (2004)
Volcanic arc of Kamchatka: a province with high d
18
O magma
sources and large-scale
18
O/
16
O depletion of the upper crust.
Geochim Cosmochim Acta 68:841–865
Blatt H (1987) Perspectives: oxygen isotopes and the origin of
quartz. J Sed Pet 57:373–377
Bleeker W (2002) Archaean tectonics: a review, with illustrations
from the Slave Craton. Geol Soc Spec Pub 199:151–181
Bowring SA, Housh T (1995) The Earth’s early evolution. Science
269:1535–1540
Burdett JW, Grotzinger JP, Arthur MA (1990) Did major changes
in the stable-isotope composition of Proterozoic seawater
occur? Geology 18:227–230
Campbell IH (2003) Constraints on continental growth models
from Nb/U ratios in the 3.5 Ga Barberton and other Archaean
basalt-komatiite suites. Am J Sci 303:319–351
Cavosie AJ, Wilde SA, Liu D, Weiblen PW, Valley JW (2004)
Internal zoning and U-Th-Pb chemistry of Jack Hills detrital
zircons: a mineral record of early Archean to Mesoproterozoic
(43481576 Ma) magmatism. Precam Res 135:251–279
Fig. 12 Plot of d
18
O(Zrc) vs. SiO
2
content for magmatic zircons
and their host rocks for 90 samples from the Neoproterozoic of
Brazil
577
Cavosie AJ, Valley JW, Wilde SA, EIMF (2005) Magmatic d
18
Oin
4400–3900 Ma detrital zircons: a record of the alteration and
recycling of crust in the Early Archean. Earth Planet Sci Lett
235:663–681
Chappell BW, White AJR (2001) Two contrasting granite types:
25 years later. Austral J Ear Sci 48:489–499
Chen D, Deloule E, Cheng H, Xia Q, Wu Y (2003) Preliminary
study of microscale zircon oxygen isotopes for Dabie-Sulu
metamorphic rocks: ion probe in situ analyses. Chin Sci Bull
48:1670–1678
Compston W, Pidgeon RT (1986) Jack Hills, evidence of more very
old detrital zircons in Western Australia. Nature 321:766–769
Condie KC (1993) Chemical composition and evolution of the
upper continental crust: contrasting results from surface sam-
ples and shales. Chem Geol 104:1–37
Condie KC (1998) Episodic continental growth and super-
continents: a mantle avalanche connection? Earth Planet Sci
Lett 163:97–108
Condie KC (2000) Episodic continental growth models: after
thoughts and extensions. Tectonophysics 322:53–162
Condie KC, Des Marais DJ, Abbott D (2001) Precambrian su-
perplumes and supercontinents: a record in black shales, carbon
isotopes, and paleoclimates? Precam Res 106:239–260
Eiler JM (2001) Oxygen isotope variations of basaltic lavas and
upper mantle rocks. In: Valley JW, Cole DR (eds) Stable iso-
tope geochemistry, Reviews in Mineralogy and Geochemistry,
vol 43. Mineralogical Society of America/Geochemical Society,
Washington, DC, pp 319–364
Eiler JM, McInnes B, Valley JW, Graham CM, Stolper EM (1998)
Oxygen isotope evidence for slab-derived fluids in the sub-arc
mantle. Nature 393:777–781
Elliott BA, Peck WH, Ramo OT, Vaasjoki M, Nironen M (2005)
Magmatic zircon oxygen isotopes of 1.881.87 Ga orogenic
and 1.651.54 Ga anorogenic magmatism in Finland. Mineral
Petrol (in press)
Eriksson KA (1995) Crustal growth, surface processes, and atmo-
spheric evolution on the early Earth. Geol Soc Spec Pub 95:11–25
Ferreira VP, Valley JW, Sial AN, Spicuzza MJ (2003) Oxygen
isotope compositions and magmatic epidote from two con-
trasting metaluminous granitoids, NE Brazil. Contrib Mineral
Petrol 145:205–216
Gilliam CE, Valley JW (1997) Low d
18
O magma, Isle of Skye,
Scotland: evidence from zircons. Geochim Cosmochim Acta
61:4975–4981
Gregory RT, Taylor HP (1981) An oxygen isotope profile in a
section of Cretaceous oceanic crust, Samail Ophiolite, Oman:
evidence for d
18
O buffering of the oceans by deep (> 5 km)
seawater-hydrothermal circulation at mid-ocean ridges. J
Geophys Res 86:2737–2755
Hanchar JM, Hoskin PWO (eds) (2003) Zircon. Reviews in Mi-
neralogy and Geochemistry, vol 53. Mineralogical Society of
America/Geochemical Society, Washington, DC, p 500
Harmon RS, Hoefs J (1995) Oxygen isotope heterogeneity of the
mantle deduced from global
18
O systematics of basalts from
different geotectonic settings. Contrib Mineral Petrol 120:95–114
Hildreth W, Christiansen RL, O’Neil JR (1984) Catastrophic iso-
topic modification of rhyolitic magma at times of caldera sub-
sidence, Yellowstone Plateau volcanic field. J Geophys Res
89:8339–8369
Hurley PM, Rand JR (1969) Pre-drift continental nuclei. Science
164:1229–1242
Kemp AIS, Hawkesworth CJ (2003) Granitic perspectives on the
generation and secular evolution of the continental crust. Treat
Geochem 3:349–410
King EM (1997) Oxygen isotope study of igneous rocks from the
Superior Province, Canada. MSc Thesis, University of Wis-
consin
King EM (2001) Oxygen isotope study of magmatic source and
alteration of granitic rocks in the western United States and the
Superior Province, Canada. PhD Thesis, University of Wis-
consin
King EM, Valley JW (2001) The source, magmatic contamination,
and alteration of the Idaho Batholith. Contrib Mineral Petrol
142:72–88
King EM, Barrie CT, Valley JW (1997) Hydrothermal alteration of
oxygen isotope ratios in quartz phenocrysts, Kidd Creek Mine,
Ontario: magmatic values are preserved in zircon. Geology
25:1079–1082
King EM, Valley JW, Davis DW, Edwards GR (1998) Oxygen
isotope ratios of Archean plutonic zircons from granite-green-
stone belts of the Superior Province: Indicator of magmatic
source. Precam Res 92:365–387
King EM, Valley JW, Davis DW (2000) Oxygen isotope evolution
of volcanic rocks at the Sturgeon Lake volcanic complex, On-
tario. Can J Earth Sci 37:39–50
King EM, Valley JW, Davis DW, Kowallis BJ (2001) Empirical
determination of oxygen isotope fractionation factors for tita-
nite with respect to zircon and quartz. Geochim Cosmochim
Acta 65:3165–3175
King EM, Valley JW, Stockli DF, Wright JE (2004) Oxygen iso-
tope trends of granitic magmatism in the Great Basin: location
of the Precambrian craton boundary as reflected in zircons.
Geol Soc Am Bull 116:451–462
Knauth LP, Lowe DR (2003) High Archean climatic temperature
inferred from oxygen isotope geochemistry of cherts in the
3.5 Ga Swaziland Supergroup, South Africa. Geol Soc Am Bull
115:566–580
Kramers JD (2002) Global modeling of continent formation and
destruction through geological time and implications for CO
2
drawdown in the Archaean Eon. Geol Soc Spec Pub 199:259–274
Kro
¨ner A, Layer PW (1992) Crust formation and plate motion in
the early Archean. Science 256:1405–1411
Lackey JS (2005) The magmatic and alteration history of the Sierra
Nevada batholith as recorded by oxygen isotope ratios in zir-
con, titanite, garnet, and quartz. PhD Thesis, University of
Wisconsin
Lackey JS, Valley JW, Saleeby JB (2005a) Supracrustal input to
magmas in the deep crust of Sierra Nevada batholith: evi-
dence from high-d
18
O zircon. Earth Planet Sci Lett 235:315–
330
Lackey JS, Valley JW, Hinke HJ (2005b) Deciphering the source
and contamination history of peraluminous magmas using d
18
O
of accessory minerals: examples from garnet-bearing granitoids
of the Sierra Nevada batholith. Contrib Mineral Petrol (in
press)
Land LS, Lynch FL (1996) d
18
O values of mudrocks: more evi-
dence for an
18
O-buffered ocean. Geochim Cosmochim Acta
60:3347–3352
Longstaffe FJ, Schwarcz HP (1977)
18
O/
16
O of Archean clastic
metasedimentary rocks: a petrogenetic indicator for Archean
gneisses? Geochim Cosmochim Acta 41:1303–1312
Lowe DR (1992) Major events in the geological development of the
Precambrian Earth. In: Schopf JW, Klein C (eds) The Proter-
ozoic biosphere a multidisciplinary study. Cambridge Uni-
versity Press, Cambridge, pp 67–75
Lowe DR (1994) Archean greenstone-related sedimentary rocks.
Dev Precam Geol 11:121–169
Lowe DR, Tice MM (2004) Geologic evidence for Archean at-
mospheric and climatic evolution: fluctuating levels of CO
2
,
CH
4
, and O
2
with an overriding tectonic control. Geology
32:493–496
Lowry D, Appel PWU, Rollinson HR (2003) Oxygen isotopes of
an early Archaean layered ultramafic body, southern West
Greenland: implications for magma source and post-intrusion
history. Precam Res 126:273–288
Lugovaya IP, Krivdik SG, Ponomarenko AN (2001) Oxygen iso-
tope composition of zircons in granites and alkaline rocks of the
Ukrainian Shield [Russian]. Mineral J 23:38–41
Marcantonio F, McNutt RH, Dickin AP, Heaman LM (1990)
Isotopic evidence for the crustal evolution of the Frontenac
Arch in the Grenville Province of Ontario, Canada. Chem Geol
83:297–314
578
McLennan SM, Taylor SR, Hemming SR (2005) Composition,
differentiation, and evolution of continental crust: constraints
from sedimentary rocks and heat flow. In: Brown M, Rushmer
T (eds) Evolution and differentiation of the continental crust.
Cambridge University Press, Cambridge, pp 93135
Mojzsis SJ, Harrison TM, Pidgeon RT (2001) Oxygen-isotope
evidence from ancient zircons for liquid water at the earth’s
surface 4,300 Myr ago. Nature 409:178–181
Monani S, Valley JW (2001) Oxygen isotope ratios of zircon:
magma genesis of low d
18
O granites from the British Tertiary
igneous province, western Scotland. Earth Planet Sci Lett
184:377–392
Muehlenbachs K (1998) The oxygen isotopic composition of the
oceans, sediments and the seafloor. Chem Geol 145:263–273
Mueller PA, Wooden JL, Nutman AP (1992) 3.96 Ga zircons from
an Archean quartzite, Beartooth Mountains, Montana. Geol-
ogy 20:327–330
Mueller PA, Wooden JL, Nutman AP, Mogk DW (1998) Early
Archean crust in the northern Wyoming Province: evidence
from U-Pb ages of detrital zircons. Precam Res 91:295–307
O’Connor YL, Morrison J (1999) Oxygen isotope constraints on the
petrogenesis of the Sybille Intrusion of the Proterozoic Laramie
anorthosite complex. Contrib Mineral Petrol 136:81–91
O’Neil JR, Chappell BW (1977) Oxygen and hydrogen isotope
relations in the Berridale batholith. J Geol Soc Lond 133:559–
571
O’Neil JR, Shaw SE, Flood RH (1977) Oxygen and hydrogen
isotope compositions as indicators of granite genesis in the New
England Batholith, Australia. Contrib Mineral Petrol 62:313–
328
Pan Y, Fleet ME (1995) Geochemistry and origin of cordierite-
orthoamphibole gneiss and associated rocks at an Archaean
volcanogenic massive sulphide camp; Manitouwadge, Ontario,
Canada. Precam Res 74:73–89
Peck WH (2000) Oxygen isotope studies of Grenville Meta-
morphism and Magmatism. PhD Thesis, University of Wis-
consin-Madison
Peck WH, Valley JW (2005) The Archean environment. In: Gornitz
V (ed) Encyclopedia of paleoclimatology and Ancient En-
vironments. Kluwer, New York (accepted)
Peck WH, King EM, Valley JW (2000) Oxygen isotope perspective
on Precambrian crustal growth and maturation. Geology
28:363–366
Peck WH, Valley JW, Wilde SA, Graham CM (2001) Oxygen
isotope ratios and rare earth elements in 3.3 to 4.4 Ga zircons:
ion microprobe evidence for high d
18
O continental crust and
oceans in the early Archean. Geochim Cosmochim Acta
65:4215–4229
Peck WH, Valley JW, Corriveau L, Davidson A, McLelland J,
Farber D (2004) Constraints on terrane boundaries and origin of
1.18 to 1.13 Ga granitoids of the Southern Grenville Province
from oxygen isotope ratios of zircon. In: Tollo RP, McLelland J,
Corriveau L, Bartholomew MJ (eds) Proterozoic evolution of the
Grenville orogen in North America, Memoir, vol 197. Geological
Society of America, Boulder, CO, pp 163–181
Perry EC, Lefticariu L (2003) Formation and geochemistry of
Precambrian cherts. Treat Geochem 7:99–113
Rumble D, Giorgis D, Ireland T, Zhang Z, Xu H, Yui T-F, Yang J,
Xu Z, Liou J-G (2002) Low d
18
O zircons, U-Pb dating, and the
age of the Qinglongshan oxygen and hydrogen isotope anomaly
near Donghai in Jiangsu Province, China. Geochim Cosmo-
chim Acta 66:2299–2306
Savin SM, Epstein S (1970) The oxygen and hydrogen isotope geo-
chemistry of clay minerals. Geochim Cosmochim Acta 34:5–42
Shieh YN (1985) High
18
O granitic plutons from the Frontenac
axis, Grenville Province of Ontario, Canada. Geochim Cos-
mochim Acta 49:117–123
Shieh YN, Schwarcz HP (1978) The oxygen isotope composition of
the surface crystalline rocks of the Canadian Shield. Can J
Earth Sci 15:1773–1782
Shields G, Veizer J (2002) Precambrian marine carbonate isotope
database: Version 1.1. Geochem Geophys Geosyst 3. DOI
10.1029/2001GC000266
Shirey SB, Hanson GN (1984) Mantle-derived Archaean mon-
zodiorites and trachyandesites. Nature 310:222–224
Simon L, Lecuyer C (2002) Continental recycling: the oxygen iso-
tope point of view. Geochim Cosmochim Acta 66:717
Stein M, Hoffman AW (1994) Mantle plumes and episodic crustal
growth. Nature 372:63–68
Stern RA, Hanson GN (1991) Archean high-Mg granodiorite: a
derivative of light rare earth element-enriched monzodiorite of
mantle origin. J Petrol 32:201–238
Sylvester PJ (1994) Archean granite plutons. Dev Precam Geol
11:261–314
Sylvester PJ (ed) (2000) Continent formation, growth and re-
cycling. Tectonophysics, vol 322. Elsevier, Amsterdam, pp 1–
202
Taylor HP (1986) Igneous rocks: II. Isotopic case studies of cir-
cumpacific magmatism. In: Valley JW, Taylor HP Jr, O’Neil JR
(eds) Stable isotopes in high temperature geological processes,
Reviews in Mineralogy, vol 16. Mineralogical Society of
America, p 273–317
Taylor SR, McLennan SM (1985) The continental crust: its com-
position and evolution. Blackwell, Edinburgh, p 312
Taylor SR, McLennan SM (1995) The geochemical evolution of the
continental crust. Rev Geophys 33:241–265
Taylor HP, Sheppard SMF (1986) Igneous rocks: I. Processes of
isotopic fractionation and isotope systematics. In: Valley JW,
Taylor HP Jr, O’Neil JR (eds) Stable isotopes in high tem-
perature geological processes, Reviews in Mineralogy, vol 16.
Mineralogical Society of America, pp 227–271
Trail D, Mojzsis SJ, Harrison TM (2005) Hadean crustal processes
revealed from oxygen isotopes and U-Th-Pb depth profiling of
pre-4 Ga detrital zircons from Western Australia. Lunar Planet
Sci Conf, Houston, XXXVI: 2223, abstract
Valley JW (2003) Oxygen isotopes in zircon. In: Hanchar JM,
Hoskin PWO (eds) Zircon, Reviews in Mineralogy and Geo-
chemistry, vol 53. Mineralogical Society of America/Geo-
chemical Society, Washington, DC, pp 343–385
Valley JW, Chiarenzelli JR, McLelland JM (1994) Oxygen isotope
geochemistry of zircon. Earth Planet Sci Lett 126:187–206
Valley JW, Kitchen N, Kohn MJ, Niendorf CR, Spicuzza MJ
(1995) UWG-2, a garnet standard for oxygen isotope ratios:
strategies for high precision and accuracy with laser heating.
Geochim Cosmochim Acta 59:5223–5231
Valley JW, Kinny PD, Schulze DJ, Spicuzza MJ (1998) Zircon
Megacrysts from Kimberlite: Oxygen Isotope Variability
Among Mantle Melts. Contrib Mineral Petrol 133:1–11
Valley JW, Peck WH, King EM, Wilde SA (2002) A cool early
Earth. Geology 30:351–354
Valley JW, Bindeman IN, Peck WH (2003) Empirical calibration of
oxygen isotope fractionation in zircon. Geochim Cosmochim
Acta 67:3257–3266
Veizer J (1983) Geologic evolution of the archean-early proterozoic
Earth. In: Schopf JW (ed) Earth’s earliest biosphere: its origin
and evolution. Princeton University Press, Princeton, NJ, pp
240–259
Veizer J, Jansen SL (1985) Basement and sedimentary recycling: 2,
Time dimension to global tectonics. J Geol 93:625–643
Veizer J, Mackenzie FT (2003) Evolution of sedimentary rocks.
Treat Geochem 7:369–407
Walker JCG, Lohmann KC (1989) Why the oxygen isotopic
composition of seawater changes with time. Geo Res Lett
16:323–326
Wallmann K (2001) The geological water cycle and the evolution
of marine d
18
O values. Geochim Cosmochim Acta 65:2469–
2485
Wei C-S, Zheng Y-F, Zhao Z-F, Valley JW (2002) Oxygen and
neodymium isotope evidence for recycling of juvenile crust in
Northeast China. Geology 30:375–378
579
Wilde SA, Valley JW, Peck WH, Graham CM (2001) Evidence
from detrital zircons for the existence of continental crust and
oceans on the Earth 4.4 Gyr ago. Nature 409:175–178
Windley BF (1995) The evolving continents. John Wiley and Sons,
Chichester, UK, p 526
de Wit MJ (1998) On Archean granites, greenstones, cratons and
tectonics: does the evidence demand a verdict? Precam Res
91:181–226
Zhao Z-F, Zheng Y-F, Wei C-S, Wu Y-B (2004) Zircon isotope
evidence for recycling of subducted continental crust in post-
collisional granitoids from the Dabie terrane in China. Geophys
Res Lett 31. DOI 10:1029/2004GL021061
Zheng Y-F, Wu Y-B, Chen F-K, Gong B, Li L, Zhao Z-F (2004)
Zircon U-Pb and oxygen isotope evidence for a large-scale
18
O
depletion event in igneous rocks during the Neoproterozoic.
Geochim Cosmochim Acta 68:4145–4165
580
