JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 9
PAGES 1901–1924
2005
doi:10.1093/petrology/egi042
Garnetite Xenoliths and Mantle–Water
Interactions Below the Colorado Plateau,
Southwestern United States
DOUGLAS SMITH1* AND WILLIAM L. GRIFFIN2,3
1
DEPARTMENT OF GEOLOGICAL SCIENCES, JOHN A. AND KATHERINE G. JACKSON SCHOOL OF GEOSCIENCES,
THE UNIVERSITY OF TEXAS AT AUSTIN, 1 UNIVERSITY STATION C-1100, AUSTIN, TX 78712-0254, USA
2
GEMOC KEY CENTRE, DEPARTMENT OF EARTH AND PLANETARY SCIENCES, MACQUARIE UNIVERSITY, SYDNEY,
NSW 2109, AUSTRALIA
3
CSIRO EXPLORATION AND MINING, NORTH RYDE, NSW 2113, AUSTRALIA
RECEIVED JULY 7, 2004; ACCEPTED MARCH 21, 2005
ADVANCE ACCESS PUBLICATION MAY 6, 2005
Garnetite xenoliths from ultramafic diatremes in northeastern
Arizona provide insights into hydration and metasomatism in the
mantle. The garnetites contain more than 95% garnet, some of
which has complex compositional zonation related to growth in
fractures within grains. Accessory minerals include rutile, ilmenite,
chlorite, clinopyroxene, and zircon. Zircon grains in one rock were
analyzed in situ to determine U–Pb ages and Hf isotopic compositions. Most U–Pb analyses plot on or near concordia in the range
60–85 Ma but a few are discordant. The range in 176Hf/177Hf is
about 02818–02828, with grains zoned to more radiogenic Hf
from interiors to rims. The garnetite protolith contained zircons at
least 18 Ga in age, and garnet and additional zircon crystallized
episodically during the interval 85–60 Ma. The garnetites are
interpreted as mantle analogues of rodingites, formed in metasomatic
reaction zones caused by water–rock interactions in Proterozoic
mantle during late Cretaceous and Cenozoic subduction of the
Farallon plate. Associated eclogite xenoliths may have been parts
of these same reaction zones. The rodingite hypothesis requires
serpentinization in the mantle wedge 700 km from the trench, beginning 5–10 Myr before tectonism related to low-angle subduction.
Garnetite xenoliths from a diatreme cluster in the Navajo
volcanic field of the Colorado Plateau have been studied
to investigate implications for serpentinization and
metasomatism of continental mantle. Helmstaedt &
Schulze (1988) suggested that these xenoliths formed by
reactions involving peridotite hydration. If so, they may
be samples of metasomatic reaction zones analogous to
those that form at low temperatures and pressures during
serpentinization (Coleman, 1967); garnet-rich parts of
such zones are called rodingites. If the garnetites are
analogues of rodingites, then they may provide evidence
of mantle processes associated with subduction of the
Farallon plate, by analogy with origins proposed for
eclogite xenoliths from the same diatremes. The locality
is about 700 km from the Farallon subduction zone
according to the reconstruction of Severinghaus &
Atwater (1990); however, this is more than three times
the maximum distance from the trench that Hyndman &
Peacock (2003) considered a typical limit for subductioninduced serpentinization. Evidence that the garnetites are
rodingite analogues would be consistent with the hypothesis of Humphreys et al. (2003) that low-angle subduction hydrated the mantle wedge below a broad region of
western North America.
The garnetites also have been studied to clarify the
genesis of the associated eclogite xenoliths. The eclogites
have been interpreted either as fragments of the Farallon
slab itself (Helmstaedt & Doig, 1975; Usui et al., 2003)
or as products of water–rock reactions in the mantle
wedge (Smith et al., 2004). These contrasting hypotheses
*Corresponding author. Telephone: either 512-471-4261 or 970-2590558. E-mail: [email protected]. (Contact Smith by e-mail before
postal mailings, because he will be at an alternate postal address for
parts of 2005. His e-mail address will remain the same.)
The Author 2005. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oupjournals.org
KEY WORDS:
garnetite; Lu–Hf, mantle; rodingite; metasomatism
INTRODUCTION
VOLUME 46
imply very different effects of low-angle subduction on
the mantle roots of continents. The garnetite xenoliths
are hosted only by those diatremes that host the eclogites,
and Switzer (1975) suggested that these rock types were
genetically related. If so, the garnetites yield insights into
how the eclogites formed.
NUMBER 9
SEPTEMBER 2005
109° 03' W
JOURNAL OF PETROLOGY
UT
500 km
CO
Colorado
Plateau
OCCURRENCE
All garnetite xenoliths studied here are from the Garnet
Ridge diatreme cluster on the Comb Ridge monocline in
northeastern Arizona (Fig. 1). The host rocks in the
diatremes are serpentinized ultramafic microbreccia
(SUM; Roden, 1981). These SUM host rocks appear to
have been emplaced as gas–solid mixes (McGetchin &
Silver, 1972), and textural and chemical evidence supports the conclusion that a melt phase was never part
of the mixture (Roden, 1981). Rather, the gas–solid
eruptions have been interpreted as products of hydrated
mantle disaggregated during heating by intruded magma
(Smith & Levy, 1976). The Garnet Ridge diatremes were
emplaced at 30 Ma (Smith et al., 2004).
Garnetite occurs in the diatremes together with an
extraordinary variety of xenoliths of sedimentary, metamorphic, and igneous rocks: gabbro, granite, rhyolite,
granulite, amphibolite, peridotite, omphacite pyroxenite,
and eclogite are among the rock types present. Eclogite
makes up less than 005 wt % of the population of igneous
plus metamorphic rocks (Hunter, 1979). The abundance
of garnetite xenoliths was not measured, but they are less
common than those of eclogite, and they probably make
up less than 001 wt % of that population. Garnetite also
occurs in the two other major diatremes on the Comb
Ridge monocline, Mule Ear and Moses Rock (Helmstaedt
& Schulze, 1988; McGetchin & Silver, 1970).
DESCRIPTIONS AND
PETROGRAPHY
The largest of the 14 garnetite xenoliths studied in thin
section had a maximum diameter of 8 cm, slightly smaller than the 11 cm dimension of the largest garnetite
xenolith described from the province by Switzer (1975).
Most were less than 4 cm in maximum diameter. Typical
specimens are nearly equant and have smooth surfaces,
some of which appear polished. These smooth surfaces
are attributed to abrasion and impact during emplacement of the gas–solid mix that formed the diatreme fill, as
described by McGetchin & Silver (1970, 1972). Irregular
fractures, some coated with films of secondary calcite,
perhaps caliche, are present in most samples.
The garnetite xenoliths consist of 95% to almost 100%
garnet. Grain boundaries are difficult to identify. Slight
variations of garnet color from clear to very pale
orange are visible in sections of some rocks. These color
37°00' N
Garnet
Ridge
Navajo
volcanic
field
AZ
NM
Fig. 1. Location of the study area within the Colorado Plateau. ~,
location of the Garnet Ridge and other major diatremes of serpentinized ultramafic microbreccia (SUM) in the Navajo volcanic field. The
boundary of the Navajo field is determined by the distribution of
minette and related rocks.
variations were seen only in polished thin sections with
thicknesses in the range from 100 to 300 mm. In some
rocks, parts of garnets are turbid with unoriented inclusions a few micrometers in average diameter, many of
which appear to be of fluid. The grain boundaries of
garnets were inferred primarily from the color variations
and turbidity and from cracks. Where estimates were
possible, garnet grains appeared anhedral and approximately equant, typically with maximum diameters of a few
millimeters.
A diverse suite of minor and trace minerals is present;
multiple polished thin sections were prepared of most
samples so as to identify as many as possible. In approximate order of decreasing abundance, these minerals are:
rutile, ilmenite, chlorite, clinopyroxene, zircon, pyrite,
phlogopite, and apatite. Rutile and ilmenite are the only
minerals other than garnet present at a modal abundance
of 1% or more; the others are present only as trace
constituents. Grains of the two oxides typically are anhedral, with maximum dimensions of several millimeters or
less. Rutile forms an estimated 5% of one sample, but it
makes up less than 2% of more typical samples and is
absent in some. The maximum abundance of ilmenite
was estimated as about 2%; most samples contain 1%,
and it was not observed in some rocks. Chlorite was
identified in six of the 14 samples, as was clinopyroxene.
Apatite, phlogopite, and pyrite were identified in two
to four rocks. Clinopyroxene, chlorite, and apatite are
slightly more common at the edges of several of the
xenoliths in which they occur. In rocks that contain
1902
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
both ilmenite and rutile, they occur as discrete grains
and as rims around one another. Chlorite is intergrown
with rutile in some occurrences and with ilmenite in some
others. No other systematic mineral associations were
noticed.
Zircon was identified in five xenoliths, but only in one
of these rocks (GR1-201) were more than a few grains
observed. Although zircon is present only in trace proportion, it is the most common silicate other than garnet
in that rock. Zircon grains in GR1-201 were chosen
for in situ U–Pb and Lu–Hf analysis. Most grains are
anhedral to subhedral, but rare grains are euhedral;
maximum diameters range from several micrometers to
about 300 mm.
MINERAL CHEMISTRY
Procedures for elemental analysis
Minerals in 11 xenoliths were analyzed by wavelengthdispersive spectrometry (WDS) on JEOL 7300 and 8200
electron microprobes at The University of Texas at
Austin. For most analyses, the accelerating voltage was
15 keV and the beam current was about 40 nA. Counting
at peak and background wavelengths was terminated at
40 s for each, unless a standard deviation <03% was
achieved first, based on counting statistics. In a few analytical sessions, trace elements were analyzed with a
20 keV accelerating voltage, higher beam current, and
longer times for data acquisition. Data were corrected
with a JEOL ZAF procedure. Representative electron
microprobe (EMP) analyses of minerals in six of the xenoliths are given in Table 1, and representative analyses for
all of the xenoliths are given in the Electronic Appendix
(available at http://www.petrology.oupjournals.org).
Trace elements in minerals of five rocks were analyzed
by laser ablation microprobe–inductively coupled plasma
mass spectroscopy (LAM–ICPMS) at The University of
Texas. After photography and characterization by EMP
analysis, minerals in polished thin sections were ablated
with a New Wave LUV213 laser. The Nd:YAG source
delivers a focused 213 nm beam; resulting ablation pits
have diameters from 50 to 100 mm. The ablated material
was transported with a helium carrier gas into a Micromass Platform quadrupole inductively coupled plasma
mass spectrometer. Signal intensities were tabulated for
each mass sweep in most of the analyses, and anomalous
data were excluded. Concentrations in garnet and
clinopyroxene were calculated with 44Ca as an internal
standard in two of the three datasets and with 43Ca in the
other: reanalysis verified consistency of the two choices.
Two semiquantitative analyses of rutile were calculated
with 93Nb as an internal standard. Four NIST SRM
glasses (610, 612, 614, and 616) and USGS glass BCR2G were analyzed during each session. For almost all
analyzed masses, calibration curves were fitted either to
SRM 616, 614, and 612 or to all four of the NIST glasses,
depending upon the relative count rate for the unknown.
Glass BCR-2G was used as a secondary standard, except
in rare cases when it was included as a primary standard
to avoid extrapolations to much higher concentrations
than those in NIST SRM 610. Consistency of some of the
garnet and pyroxene analyses was tested and verified
by analysis of two masses for each of five REE. Accuracy
was verified by the analyses of the secondary standard,
BCR-2G. Backgrounds were obtained by measurements
without ablation of samples. Detection limits were not
calculated from these backgrounds, because of possible
differences between signal levels with and without sample
ablation. Instead, results are not reported below determination limits based on average concentrations within
two standard deviations of zero. Representative determination limits were about 55 ppb for La in garnet and
near and below 12 ppb for Tm, Yb, and Lu in pyroxene.
Garnet
The most common garnet compositions have roughly
equal atomic proportions of Mg, Fe, and Ca (Fig. 2),
consistent with the more limited data of McGetchin &
Silver (1970), Switzer (1975) and Helmstaedt & Schulze
(1988). The compositions average about 2% andradite,
calculated from (2 – Al – Cr)/2 and 8-cation formulae.
Other minor elements are present with the following
average contents: TiO2, 007 wt %; Cr2O3, 005 wt %;
MnO, 034 wt %; Na2O, 003 wt %. No clear correlations were observed between major (Ca, Fe, Mg) and
minor elements (Na, Ti, Cr, and Mn) in the entire database, although Fe and Mn are well correlated in the
garnet of some rocks. The most magnesian garnets
(50% pyrope end-member) are found in xenolith N379bGR-4, and they have more Cr2O3 (013–017 wt %) than
garnet in other samples.
Many garnets are complexly zoned, as illustrated in
back-scattered electron (BSE) images in Fig. 3. Two types
of zoning were distinguished. First, gradients from grain
interiors to rims were observed (Fig. 3a and b). Second,
compositional gradients also are present in patterns
that record both growth of garnet within fractures and
reaction of the fractured grains (Fig. 3c). The more complex patterns formed by the interplay of both processes
(Fig. 3d and e).
No trends of compositional zoning are consistent for
the full dataset, although trends in individual garnets are
well defined (Figs 4 and 5). BSE images document compositional changes at rims adjacent to ilmenite in rock
N375-GR (Fig. 3a and b): one such rim, about 400 mm
wide, is Mg-rich and Ti-poor relative to the interior
(Fig. 4a). Annular orange zones are present in garnets of
rock GR1-202, and a traverse across such a zone and an
1903
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 9
SEPTEMBER 2005
Table 1: Representative electron microprobe analyses of minerals
Sample:
GR1-203
Mineral:
Gar1
SiO2
TiO2
40.2
0.05
Chl1
Chl
30.8
0.01
31.2
0.07
19.0
0.02
Al2O3
22.0
19.3
Cr2O3
n.a.
16.6
n.a.
4.86
FeOt
MnO
MgO
CaO
Na2O
NiO
Total
0.29
7.62
0.00
31.1
14.0
0.00
0.03
0.01
n.a.
100.8
Sample:
N379b-GR-4
Mineral:
Gar1
SiO2
TiO2
Al2O3
Cr2O3
FeOt
MnO
MgO
CaO
Na2O
NiO
Nb2O5
Total
18.9
0.10
14.4
0.35
3.48
0.02
14.0
8.11
32.6
0.02
0.00
0.02
0.02
0.37
Mineral:
Gar1
TiO2
Al2O3
Cr2O3
FeOt
MnO
MgO
CaO
Na2O
NiO
Total
31.6
0.01
0.01
0.41
Chl1
23.2
0.16
ATG-GRP1-B
4.87
0.02
Gar3
Cpx3
Rut
Ilm
40.2
0.10
55.9
0.07
39.6
0.10
55.4
0.04
0.06
96.8
0.06
51.5
22.1
5.86
0.01
3.41
21.8
0.04
2.78
0.03
0.22
0.11
0.06
0.03
18.0
0.29
3.70
0.02
2.58
0.00
0.24
8.56
0.01
12.5
12.8
0.03
18.2
4.41
n.a.
100.7
8.48
11.6
0.02
14.3
21.5
2.42
0.07
0.02
100.1
n.a.
100.4
100.2
43.3
0.1
0.02
0.01
4.62
0.00
0.00
0.00
0.00
0.20
99.8
99.8
GR1-201
32.0
0.00
n.a.
101.8
Cpx2
n.a.
16.7
87.3
41.3
0.05
Sample:
SiO2
n.a.
86.1
Gar2
Rut4
Rut5
Gar
0.02
53.8
0.05
n.a.
93.4
0.18
n.a.
94.9
39.4
0.07
39.2
0.02
0.06
39.3
0.63
2.00
21.8
0.09
22.1
0.02
17.7
0.49
23.1
1.00
7.91
11.9
8.04
6.90
n.a.
0.03
0.00
0.04
0.00
n.a.
Ilm
0.14
7.10
n.a.
n.a.
n.a.
n.a.
0.00
0.01
n.a.
n.a.
n.a.
n.a.
n.a.
1.20
n.a.
0.62
97.4
98.1
0.21
n.a.
87.5
0.17
0.53
1.82
n.a.
100.7
Gar
n.a.
99.4
Rut
n.a.
100.5
N379b-GR-1
Chl1
Gar2
Cpx2
Ilm
Rut
Gar6
n.a.
96.5
0.18
0.24
1.82
n.a.
n.a.
n.a.
0.04
98.8
GR1-202
Gar7
Gar
40.5
0.08
32.6
0.00
40.0
0.05
54.2
0.00
0.00
54.1
0.02
97.2
39.2
0.09
39.8
0.07
39.8
0.10
22.8
0.01
18.9
0.08
22.9
0.05
22.3
0.02
15.0
0.31
0.17
0.29
2.48
22.5
0.07
3.51
0.01
0.07
0.08
40.6
21.5
0.02
15.4
0.28
0.81
0.02
2.08
16.0
0.24
12.3
10.3
0.02
0.01
12.9
0.14
31.9
0.02
0.09
6.58
18.8
0.25
12.5
9.83
0.02
17.1
8.88
0.02
0.00
0.01
0.28
0.02
0.00
0.04
0.01
6.46
13.6
0.05
8.33
24.7
0.60
16.4
0.01
12.9
0.02
0.04
99.6
0.22
101.7
101.4
87.3
0.03
0.00
100.9
FeOt, total Fe as FeO; n.a., not analyzed.
1
Pairs used for garnetchlorite temperature calculations.
2,3
Pairs used for garnetpyroxene temperature calculations.
4,5
Analyses of rutile with highest and lowest Nb in rock.
6,7
Interior and rim, respectively.
1904
0.00
100.2
n.a.
100.0
n.a.
100.1
n.a.
100.2
50
40
60
60
30
10
0
50
50
40
0
0
40
20
GARNETITE XENOLITHS, COLORADO PLATEAU
40
30
50
20
0
10
SMITH AND GRIFFIN
10
10
30
30
Ca
20
20
20
20
30
30
10
10
(a) Rock (# analyses)
GR1-202 (55)
GR1-203 (8)
GR-P1-ATG (7)
GR-41 (2)
0
40
Fe
50
Mg
60
40
50
60
0
(b) Rock (# analyses)
N375-GR (14)
N379b-GR-3 (4)
N379b-GR-4 (5)
ATG-GRP1-B (6)
GR1-201 (29)
N379b-GR-1 (51)
N379b-GR-2 (39)
Fig. 2. Fe–Ca–Mg proportions in garnets of 11 garnetite xenoliths. (a) Analyses of garnet in eight xenoliths. Garnets in these samples either
appeared relatively homogeneous for these elements, or too few analyses were made to determine if heterogeneities are present. (b) Analyses of
garnet in three xenoliths in which heterogeneity for these elements is well documented.
apparent grain boundary is plotted in Fig. 4b; Ti is high
at the boundary relative to the grain interior, unlike the
example plotted in Fig. 4a. The gradients in the garnet of
Fig. 5 are due to both ‘normal’ and fracture-related
growth. In a clear case of fracture-related zonation in
N379b-GR-2, the fracture fill is less calcic and less titaniferous than the surrounding garnet (Fig. 5a). Garnet in
N379b-GR-1 has complex zoning that includes both
interior-to-rim gradients and gradients along sets of
small intersecting fractures (Fig. 5b).
Trace elements determined by LAM–ICPMS in garnet
include Ni, Y, Zr, Yb, and Hf and all the rare earth
elements (REE) except La (Table 2). Measurements of
La, Nb, Ta, and U yielded concentrations below determination limits. Average Ni concentrations fall in the
range 8–18 ppm. Of garnet in the five xenoliths, that in
zircon-bearing GR1-201 is highest in Y (67 ppm), Lu
(13 ppm) and the other heavy REE (HREE), and lowest
in Zr (2 ppm) and Hf (004 ppm). Chondrite-normalized
profiles of REE abundance in garnet in each of the five
rocks are relatively similar for the range Ce to Sm
(Fig. 6a). In three of the rocks, however, normalized
REE abundances decrease with increasing atomic number in the range Sm to Lu, forming ‘humped’ patterns. In
one of these three rocks, there is a distinct positive Eu
anomaly, and in the other two, poorly defined negative
anomalies. The patterns of average REE abundance for
each rock adequately represent all the analyses of garnet
in that sample, except for garnetite N379b-GR-1.
Analyses of five of the six ablated volumes of garnet in
this rock form a clustered group, whereas the sixth is
much lower in HREE (Fig. 6b). The volume of this
anomalous analysis is distinguished by lower average
atomic weight and so is relatively dark on the BSE
image (Fig. 3d); the volume also has relatively lower Fe
and higher Mg, Ca, and Cr (Fig. 5b). Although no
comparable inhomogeneity in REE was documented in
garnet of the other four samples, analytically significant
differences in concentration are present (Fig. 6b and c).
Rutile, ilmenite, clinopyroxene, chlorite,
and phlogopite
Compositions of rutile and ilmenite are similar in all but
xenolith N379b-GR-4 (Table 1 and the Electronic
Appendix). Ilmenite, analyzed by EMP in five xenoliths,
contains 009–027 wt % NiO and 40–71 wt % MgO
and trace Cr, Al, and Mn; that in sample N379b-GR-4
has the highest Mg, Ni, and Cr. EMP analyses of rutile in
seven xenoliths establish the typical range of 01–03 wt %
Cr2O3, but rutile in N379b-GR-4 contains 05–06 wt %
Cr2O3. In two of the three rocks in which rutile was
analyzed for Nb2O5, concentrations are near 004 wt %,
but they are much higher, 062–120 wt % Nb2O5, in
N379b-GR-4. Concentrations of V and Zr in rutile,
determined by semiquantitative LAM–ICPMS analysis,
1905
VOLUME 46
(a)
NUMBER 9
SEPTEMBER 2005
(b)
(c)
Ilm
e
nit
e
JOURNAL OF PETROLOGY
100 µm
400 µm
(e)
100 µm
(d)
100 µm
1000 µm
Fig. 3. Back-scattered electron (BSE) images of garnets. Arrows mark the ends of the EMP traverses plotted in Figs 4 and 5. (a) Garnet zoning at
grain boundary with ilmenite in rock N375-GR. Diamonds mark positions of the analyses plotted in Fig. 4a. The outlined rectangle is shown with
enhanced contrast in (b). (b) Contrast-enhanced portion of field of view in (a), showing irregular compositional boundaries. (c) Apparent fracturerelated compositional zoning in rock N379b-GR-2; the path of the compositional traverse (Fig. 5a) is visible because of oil condensation under the
electron beam. (d) Complex zonation in rock N379b-GR-1. Garnet is the only mineral in this image. Diamonds show positions of analyses plotted
in Fig. 5b. The distinctive REE profile in Fig. 6b represents an analysis of material ablated from a pit near the lower right end of the EMP
traverse, within the darker (lower average atomic weight) garnet in this BSE image. (e) Fracture-related and oscillatory zoning in an enlargement
of the rectangle outlined in (d).
are about 4500 and 170 ppm, respectively, in xenolith
N379b-GR-4 and about 3000 and 60 ppm in the relatively zircon-rich rock, GR1-201.
Clinopyroxene was analyzed in five xenoliths (Tables 1
and 2, and Electronic Appendix). In three of the five, only
one small grain was found, and in the fourth, only two
grains. In three xenoliths, the Jd component of the pyroxene is in the range 3–6%, and in one it is 14%; in the
xenolith with two identified grains, one has 12% Jd and
the other 25%. Aegirine, calculated as (Na – Al – Cr) on a
four-cation basis, ranges from 05% to 6%. The dominant end-member is diopside, and calculated Fe2þ/(Fe2þ þ
Mg) of these pyroxenes ranges from 005 to 008. A relatively large grain of clinopyroxene, about 3 mm in maximum diameter, was found in one rock (ATG-GRP1-B),
together with four much smaller grains. The chondritenormalized pattern of REE abundances in the large grain
has a maximum at Nd and very low values for the HREE
(Fig. 6c). Concentrations of Ni and Sr are about 400 ppm;
those of other trace elements are much lower.
1906
SMITH AND GRIFFIN
20
GARNETITE XENOLITHS, COLORADO PLATEAU
(a)
GB
(b)
CaO
20
± 2σ
Fe as
FeO MgO
16
pale
orange
clear
CaO
CaO
40
0.4
Weight percent oxide
Weight percent oxide
12
MgO
Cr2O3
0.5
MnO
100
200
300
400
0.3
±2σ
Fe as
FeO
16
FeO
0.6
8
pale
orange
500
TiO2
12
MgO
8
GB?
4
0.4
0.3
MnO
0.2
MnO
0.1
0.1
TiO2
Cr2O3
0.0
0
100 200 300 400 500
Distance (µm) from grain boundary (GB)
0.0
0
0.2
TiO2
Cr2O3
1000 2000 3000 4000 5000
Distance (µm)
Fig. 4. Compositional zoning documented by EMP analyses. The 2s error bars show the range of 4 SD for a representative analysis, calculated
solely on the basis of counting statistics. (a) Apparent growth zoning in garnet at a contact with ilmenite in rock N375-GR, along a path marked
on the BSE image in Fig. 3a. The grain boundary is marked ‘GB’. (b) Zonation across an area with faint boundaries between a clear central area
and a pale orange halo in rock GR1-202. A probable grain boundary is marked by ‘GB?’.
Chlorite, analyzed by EMP in five xenoliths, is
clinochlore, with Fe/(Fe þ Mg) in the range 005–008
(Table 1 and Electronic Appendix). NiO contents range
from 024 to 044 wt % and average 034 wt %. Fe/Mg of
chlorite inclusions varies systematically with that of the
surrounding garnet (Fig. 7). Phlogopite was identified in
two rocks and analyzed in one. The phlogopite has
Fe/(Fe þ Mg) ¼ 010 and 022 wt % NiO.
ZIRCON GEOCHRONOLOGY
Analytical techniques
Zircons in four polished sections of garnetite GR1-201
were analyzed at GEMOC in Macquarie University.
U–Pb ages were determined by LAM–ICPMS as
described by Belousova et al. (2001) and Jackson et al.
(2004). The instrumentation consisted of a modified
Merchantek/New Wave LUV 213 nm Nd:YAG laser
attached to an Agilent 4500 s ICPMS system. Typical
ablation pits are 30 mm in diameter, and all ablations
were performed in a He carrier gas. Mass bias and
instrumental drift were corrected using a very homogeneous external standard, the GEMOC GJ-1 zircon
(608 Ma). Data were reduced using the in-house GLITTER software (www.es.mq.edu.au/GEMOC), which
allows online selection of the most stable part of the
time-resolved signal. U and Th contents of the ablated
volumes were estimated by comparison of count rates
with those of the GJ-1 standard, and probably are
accurate to 10%.
Hafnium isotopes were analyzed with a Merchantek/
New Wave LUV 213 nm LAM, attached to a NuPlasma
multicollector (MC) ICPMS system; ablations were performed in He. Typical ablation pits are 50–60 mm in
diameter. Procedures for the correction of isotopic interferences of 176Lu and 176Yb on 176Hf and for mass bias
corrections have been described by Griffin et al. (2000).
Interpretations of Hf isotope evolution utilize the 176Lu
decay constant of Blichert-Toft et al. (1997) and the model
for the depleted-mantle source adopted by Griffin et al.
(2000). U–Pb and Hf isotope data on the secondary
standards analyzed as unknowns with each run are
summarized in Table 3.
U–Pb zircon ages
All but four of the 27 analyses of 206Pb/238U and
207
Pb/235U in zircons yield ages concordant within
analytical uncertainties in the range from about 85 to
60 Ma (Fig. 8, Table 4). The 206Pb/238U ages cluster in
two ranges, the larger cluster centered at about 70 Ma,
1907
JOURNAL OF PETROLOGY
20
Filled
fracture
VOLUME 46
NUMBER 9
(b)
20
(a)
CaO
± 2σ
CaO MgO
± 2σ
Fe as
FeO
16
16
FeO
Weight percent oxide
Weight percent oxide
Fe as
FeO
12
12
MgO
CaO
8
4
0.4
MnO
0.3
MgO
8
0.6
0.5
4
300
600
900 1200 1500
0.4
TiO2
0.3
0.2
0.2
0.1
SEPTEMBER 2005
MnO
TiO2
0.1
TiO2
Cr2O3
0.0
0
50
100
150
Distance (µm)
0.0
0
200
300
600 900 1200 1500
Distance (µm)
Fig. 5. Compositional zoning documented by EMP analyses. The 2s error bars show the range of 4 SD for a representative analysis, calculated
solely on the basis of counting statistics. (a) Zoning related to apparent fracture fill and reaction in rock N379b-GR-2 along a path marked on the
BSE image in Fig. 3c. The traverse increment was 6 mm. (b) Complex zoning apparently related both to rimward growth and to fracture fill plus
reaction in rock N379b-GR-1 along a path marked on the BSE image in Fig. 3d. The steep compositional gradients at the left edge of the figure
were defined by steps at smaller increments (6 mm) than the 36 mm steps used to span the remainder of the traverse.
and a smaller one centered at about 85 Ma (Fig. 8b).
These U–Pb analyses are compatible with a model of
episodic zircon growth extending over the period from
about 85 to about 60 Ma. The most discordant point
has a 207Pb/206Pb age of about 670 Ma. The discordant
points document inheritance of a relatively small mass of
zircon of Proterozoic but otherwise poorly defined age.
The analyzed zircons show a wide range in their
contents of Th (3–237 ppm, mean 87 ppm) and U (35–
436 ppm, mean 162 ppm). About one-third of the grains
have Th/U 03 (median 049), which might be expected in metamorphic zircons, but others range up to >2.
There is no correlation between age and Th/U, but the
two oldest grains also have the highest Th contents and
the highest Th/U. These older ages may reflect incomplete Pb loss from originally magmatic grains inherited
from the protolith of the garnetite. The high mean contents of U and Th are unusual in metamorphic zircons.
Hf isotope ratios and age implications
The analyzed zircons have extraordinarily low
176
Lu/177Hf and 176Yb/177Hf ratios, comparable with
or lower than those found in kimberlitic zircons (Griffin
et al., 2000). The depletion of the HREE is consistent with
growth in the garnetite matrix, where HREE are strongly
partitioned into garnet, but Hf partitions into zircon.
There is no correlation between 176Hf/177Hf and Yb/Hf
or Lu/Hf ratios, confirming the efficacy of the isobaricoverlap corrections employed here and described by
Griffin et al. (2000).
The 176Hf/177Hf determinations extend over the range
0281829(9)–0282851(10) (Figs 9 and 10, Table 5). Hf
isotope ratios in interiors and rims of selected grains were
obtained both by multiple analyses and by ablating
through zircons. There is a crude correlation of
176
Hf/177Hf with percent HfO2; the zircons with least
radiogenic Hf values have a median of about 21 wt %
HfO2, and the most radiogenic a median of about
17 wt %. Also, there is a crude correlation of more
radiogenic Hf with younger 206Pb/238U ages (Fig. 9b).
Furthermore, in all cases, grain rims have higher
176
Hf/177Hf than interiors, and a large part of the total
range is present within single zircon grains (Fig. 10).
Some of the grains were characterized by BSE and cathodoluminescence (CL) before analysis, and representative
1908
1909
1.5 (1.31.8)
0.19 (0.160.22)
1.5 (1.31.8)
0.22 (0.170.28)
0.15 (0.120.20)
3.4 (1.64.3)
0.66 (0.290.84)
1.6 (0.72.1)
0.19 (0.100.25)
1.5 (0.72.0)
0.20 (0.090.26)
0.28 (0.100.55)
0.10 (0.070.14)
0.010 (0.010.01)
0.015 (0.010.02)
Ho
Er
b.d.l.
b.d.l.
0.13 (0.100.14)
Yb
Lu
b.d.l., below determination limit, as discussed in text.
*,yInterior and rim, respectively, of garnet (Figs 3d and 5b, and Table 1).
Hf
b.d.l.
Tm
Dy
Tb
2.7 (2.33.2)
0.59 (0.500.67)
3.0 (2.63.3)
0.46 (0.410.53)
4.9 (2.86.5)
0.64 (0.310.88)
Gd
Eu
Sm
Nd
0.59 (0.430.69)
0.70 (0.620.81)
0.05 (0.030.05)
b.d.l.
0.41 (0.200.67)
0.25 (0.160.39)
1.8)
2.7)
0.10)
0.09)
0.21 (0.19, 0.23)
3.2 (3.1, 3.3)
0.41 (0.38, 0.44)
2.6 (2.4, 2.7)
0.40 (0.36, 0.44)
4.2 (4.1, 4.4)
0.87 (0.81, 0.92)
1.4)
.
.
4 5 (4 5, 4.5)
0.65 (0.64, 0.67)
1.6 (1.5,
2.6 (2.5,
1.3 (1.3,
b.d.l.
0.09 (0.07,
0.08 (0.07,
24 (23, 24)
11.8 (10.5, 13.1)
16.8 (15.019.6)
12.5 (10.813.8)
3.7 (3.15.5)
4.9 (4.35.8)
0.88 (0.651.13)
b.d.l.
0.17 (0.120.23)
0.14 (0.100.17)
17 (924)
18 (1022)
8 (7, 9)
0.10 (0.07, 0.12)
370
2
Gar
N379b-GR-4
16 (1127)
0.84 (0.301.62)
420 (190660)
8
Gar
GR1-202
2.6 (2.13.3)
5.8 (2.17.5)
1.2 (0.91.6)
Pr
Ce
La
1.6 (1.21.8)
10.0 (8.611.9)
1.8 (1.72.0)
1.2 (1.01.5)
1.5 (1.21.9)
6.6 (4.99.0)
Zr
Y
380 (340450)
0.21 (0.190.23)
Sr
11 (315)
0.25 (0.100.34)
450 (340550)
140 (100180)
430 (330570)
Ti
Points:
Ni
Cpx
6
Mineral:
Gar
9
ATG-GRP1-B
Sample:
Table 2: Averages and ranges (ppm) of LAM–ICPMS analyses of garnet and pyroxene
10.4 (7.015.8)
1.3 (0.82.2)
0.04 (0.010.06)
2.7 (2.14.4)
8.5 (5.913.4)
1.3 (0.82.4)
6.9 (4.510.1)
1.4 (1.02.2)
11.5 (7.617.2)
0.9 (0.31.7)
3.3 (1.75.6)
1.6 (0.92.4)
b.d.l.
0.10 (0.030.26)
0.04 (0.010.07)
67 (5293)
2.3 (1.02.7)
8 (411)
0.26 (0.040.44)
270 (220350)
9
Gar
GR1-201
Gary
0.15 (0.110.19)
0.03 (0.010.05)
0.16 (0.130.18)
1.4 (1.11.6)
0.44 (0.360.49)
1.2 (1.01.6)
0.05
0.30
0.04
0.30
0.04
0.80
0.13
1.5
0.14
2.6
1.3
1.9 (1.62.3)
1.5 (1.21.7)
2.4 (2.12.8)
0.40 (0.350.47)
2.5 (2.12.9)
0.07
1.6
b.d.l.
0.11
3
3.4
0.30
17
450
1
0.05 (0.040.07)
1.2 (0.91.4)
b.d.l.
0.07 (0.060.10)
12 (1014)
3.3 (2.24.2)
18 (1523)
0.50 (0.21.0)
480 (380690)
5
Gar*
N379b-GR-1
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
VOLUME 46
100
Chondrite-normalized REE abundances
0.1
57 58 59 60
800
°C
7
0°
0
0
°C
C
60
Fe/Mg chlorite
1
0.08
0°
C
2
3
4
5
10
SEPTEMBER 2005
0.12
1
(a)
NUMBER 9
0.04
50
JOURNAL OF PETROLOGY
40
Pair in garnetite
Pair in garnet
pyroxenite
62 63 64 65 66 67 68 69 70 71
100
(b)
0
10
0.4
0.8
Fe/Mg garnet
1.2
Fig. 7. Fe/Mg in garnet and chlorite in Navajo xenoliths in textures
consistent with equilibrium. Compositions of the five pairs in garnetites
are given in the Electronic Appendix. The three pairs in garnet
pyroxenites are from Helmstaedt & Schulze (1988) and Smith (1995).
The lines of constant temperature were calculated using the geothermometer of Dickenson & Hewitt (1986) in the modified form cited by
Laird (1988).
1
0.1
57 58 59 60
100
C
0°
62 63 64 65 66 67 68 69 70 71
(c)
Garnet
10
1
Clinopyroxene
0.1
57 58 59 60
La
Ce
Pr
Nd
62 63 64 65 66 67 68 69 70 71
Sm
Eu Tb Ho Tm Lu
Gd
Dy Er
Yb
Fig. 6. Rare earth element abundances normalized to the C1 chondrite
averages of Sun & McDonough (1989). (a) Average abundances in
garnet of five rocks: 1, GR1-201, the relatively zircon-rich rock (n ¼ 9);
2, N379b-GR-4 (n ¼ 2); 3, GR1-202 (n ¼ 8); 4, ATG-GRP1-B (n ¼ 9);
5, N379b-GR-1 (n ¼ 6). (b) Analyses of six ablated volumes in rock
N379b-GR-1, all within the complexly zoned garnet imaged by BSE in
Fig. 3d. Five of the analyses cluster closely, as is typical of the analyses
for each of the other rocks plotted in (a). Analyses of the sixth volume
are markedly lower in the HREE; that volume is garnet that has a
relatively low average atomic number and so is darkest in the BSE
image (Fig. 3d), and that is relatively lower in Fe (Fig. 5b). (c) Abundances in clinopyroxene (dashed lines, six ablations) and garnet
(continuous lines, nine ablations) in rock ATG-GRP1-B. The two
curves for garnet with slightly lower Sm abundances were ablated on
a later date than the other seven, and the lower abundances of the
MREE recorded by the two might be due to either an analytical
problem or garnet inhomogeneity.
images with positions of ablated pits are shown in Fig. 11.
Most of the analyzed grains are relatively large and have
smooth shapes (Fig. 11a–c); the interiors of these grains
have relatively old U–Pb ages and relatively low
176
Hf/177Hf, consistent with the presence of an inherited
Proterozoic component in the lased volumes. Zircons
are also present that have botryoidal shapes and granular
surfaces (Fig. 11d and e). Only two grains with this
unusual morphology were large enough to analyze, and
the ablated volumes in each had relatively high
176
Hf/177Hf (02825). Zircons with similar granular
surfaces were noted in two of the fractions from Navajo
eclogites used for U–Pb geochronology by Smith et al.
(2004): neither of those multigrain fractions recorded
inheritance of Proterozoic lead, in contrast to many of
the other analyzed fractions from the eclogites. The
botryoidal morphology may characterize some of the
zircon formed in the Cenozoic metamorphic events.
The least radiogenic 176Hf/177Hf values yield depletedmantle model ages of about 19 Ga (Fig. 9), and the age of
the garnetite protolith is likely to have been at least
18 Ga. The depleted-mantle line in Fig. 9 reproduces
average mid-ocean ridge basalt (MORB) values (Griffin
et al., 2000): the range for basalt on the East Pacific Ridge
(Chauvel & Blichert-Toft, 2001), plausibly from mantle
sources like those for Farallon MORB, is indicated by
the EPR bracket in Fig. 9b. All values of 176Hf/177Hf in
the zircons are less radiogenic than those expected for
the Mesozoic and Cenozoic basalts of the Farallon plate.
1910
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
Table 3: U–Pb and Lu–Hf analyses* of standard zircons and solution measured during a period encompassing
that of this study
Sample
No. of analyses
UPb ages (Ma) of standardsy
Pb/206Pb
91500
LAMICPMS
83
TIMS
Mud
LAMICPMS
Tank
TIMS
Sample
207
2s
207
2s
206
2s
208
1069.3
3.1
1064.3
3.3
1061.0
4.0
1049
1065
735.5
1
4.7
731.2
3.2
730.6
3.9
732
5
73
Pb/235U
No. of analyses
Hf/
2s
177
Hf
solution
91500
LAMMC-
13
634
0.282165
0.282310
34
1477
0.282290
0.282522
72
Pb/232Th
2s
16
732.4
8.3
Analyses of Hf standardsy
176
JMC475
Pb/238U
40
178
Hf/
177
Hf
2s
180
Hf/
2s
177
1.467262
1.467043
53
1.467140
1.467127
24
84
1.886774
1.886809
176
Yb/
2s
177
Hf
Hf
176
Lu/
177
Hf
2s
160
222
0.01186
166
0.00509
16
0.00032
1
1
8
0.00028
0.00011
ICPMS
TIMS
Mud
LAMMC-
Tank
ICPMS
28
68
1.886797
1
*Error-weighted means and 2s errors on error-weighted means (Ludwig, 2001).
yTIMS reference values for standard zircons: 91500, Wiedenbeck et al. (1995); Mud Tank UPb, Black & Gulson (1978).
The more radiogenic 176Hf/177Hf values can be produced by episodic zircon growth in an environment
with high Lu/Hf. The eight analyses of garnet in sample
GR1-201 have a median Lu/Hf equal to 33, far higher
than the values of 1–2 characteristic of garnet in the other
garnetites. The calculated evolution of 176Hf/177Hf in an
environment with Lu/Hf of 30 and with an initial value
like that of the inherited zircon nicely fits the observed
correlation with 206Pb/238U ages during the period from
85 to 70 Ma (Fig. 9b).
Hf-isotope composition has a weak negative correlation
with U: the grains with the most radiogenic Hf (lowest
TDM, where DM indicates depleted mantle) have the
lowest U contents. Most low-Th grains also have low
TDM. There is considerable scatter in Th and U contents,
however, and no correlation was recognized between Hf
isotope composition and Th/U.
PETROLOGICAL FRAMEWORK
AND CONSTRAINTS
Processes of formation of other garnetites
Garnetites, rocks that consist mostly of garnet, have
been attributed to a variety of processes. For instance,
monomineralic garnet xenoliths in African kimberlite
have been interpreted as magmatic cumulates; the garnets contain about 70% pyrope and 10–20% grossular
end-members (Exley et al., 1983). Layers of garnetite
consisting of 60–90 vol. % garnet occur in high-grade
amphibolite- and granulite-facies rocks in the Alps, and
they have been attributed to anatexis and melt–rock
interactions (Rivalenti et al., 1997); garnets in the most
garnet-rich layers contain about 50% almandine and
10% grossular end-members. Garnetite with about
20 vol. % corundum occurs in a garnet peridotite lens
in the Sulu ultrahigh-pressure terrane and has been interpreted as a metamorphic rock derived from a protolith
of spinel websterite (Zhang et al., 2004). Grossular-rich
garnetite boudins in migmatite have been attributed to
ultrahigh-pressure metamorphism of a subducted calcsilicate protolith (Vrana & Fryda, 2003).
The most analogous garnet-rich rocks may be rodingites. Rodingites are formed in metasomatic reaction
zones during serpentinization of peridotite in lowtemperature near-surface environments, and they occur
near contacts between peridotite and a variety of other
lithologies (Coleman, 1967; Leach & Rodgers, 1978;
Schandl et al., 1989; O’Hanley et al., 1992; Dubinska
et al., 2004). Lenses of rodingite typically are less than
1911
JOURNAL OF PETROLOGY
(a)
VOLUME 46
206Pb / 238U
140
120
0.016
100
80
0.012
60
0.008
40
0.004
0.00
0.04
0.08
0.12
0.16
0.20
0.24
207Pb / 235U
Relative probability
(b)
50
70
90
110
SEPTEMBER 2005
Temperatures and pressures recorded
by Navajo garnetites
160
0.024
0.020
NUMBER 9
130
150
206Pb / 238U age (Ma)
Fig. 8. U–Pb data and ages obtained by LA–ICP-MS analysis of
zircons in garnetite GR1-201. (a) Ratios and concordia plotted using
methods of Ludwig (2001). Data points are represented by 1s error
ellipses. (b) Probability density diagram illustrating the uneven distribution of 206Pb/238U ages.
30 cm wide and are rarely more than 5 m in maximum
dimension. The most common garnet in rodingite is finegrained hydrogrossular; andradite is much less common.
Most rodingites contain other minerals, such as diopside,
prehnite, zoisite, and tremolite. Temperatures calculated
from the mineral assemblages typically are in the range
200–500 C, but the higher temperatures may record
metamorphism after rodingite formation, as described
by Rice (1983). Some instances of the recorded metamorphism occurred at crustal pressures (e.g. Frost, 1975)
and others at mantle pressures (e.g. Evans et al., 1979).
A metarodingite–eclogite suite at Cima di Gagnone in
the Swiss Alps may be of particular pertinence to the
Navajo garnetites, because these Alpine rocks have been
interpreted to record maximum pressures near 25 GPa
(Evans et al., 1979, 1981). Garnets in the most calcic
lenses at Cima di Gagnone have grossularite components
exceeding 75%, but compositions of garnets in rocks
transitional between eclogites and metarodingites are
less calcic and overlap the range present in the Navajo
garnetites.
Calculations of pressures and temperatures recorded by
the garnetite xenoliths depend upon assumptions that
mineral grains in contact had been in equilibrium. The
ranges of garnet composition within some of the small
xenoliths and the sharp compositional boundaries within
some grains (Fig. 3) are evidence of disequilibrium, as is
the contrast in jadeite component (12 and 25%) between
the only two clinopyroxene grains found in garnetite
GR1-203 (Table 1). None the less, an approach to
equilibrium is consistent with interelement correlations
between samples. For instance, in xenolith N379b-GR-4,
garnet, chlorite and ilmenite have the most magnesian
and chrome-rich compositions within the suite. The systematic Fe–Mg partitioning between garnet and chlorite
grains (Fig. 7) is consistent with equilibrium over a temperature range of about 100 C. The partitioning of REE
between the single garnet–clinopyroxene pair analyzed
for trace elements provides another test. The chondritenormalized abundances of REE in clinopyroxene form a
bell-shaped pattern (Fig. 6c). Both the abundance pattern
and the relative depletion of HREE in the clinopyroxene
are consistent with equilibration with garnet (e.g. Harte &
Kirkley, 1997). Because the P–T and compositional
dependences of trace element partitioning between
garnet and clinopyroxene are poorly known (Blundy &
Wood, 2003), quantitative evaluation of equilibrium
using the partitioning is not yet possible for the mineral
pair in this xenolith.
Compositions used for temperature calculation were
acquired near mutual contacts of chlorite and garnet
and of clinopyroxene and garnet. Temperatures were
calculated from Fe/Mg partitioning for chlorite–garnet
pairs and Fe2þ/Mg partitioning for clinopyroxene–
garnet pairs at a pressure of 2 GPa (Table 6); no pressures
have been calculated from the mineral assemblages in
these xenoliths. The thermometer of Krogh (1988) and
the revised thermometer of Krogh Ravna (2000) have
been applied to garnet–pyroxene pairs; for these applications, ferric iron was calculated in pyroxene from
(Na – Al – Cr) in a four-cation formula, and in garnet
from (2 – Al – Cr) in an eight-cation formula. Temperatures calculated with the more recent procedure, that of
Krogh Ravna (2000), range between 460 C and 600 C
and average 530 C. Values calculated with the procedure of Krogh (1988) are 60–90 higher and average about
600 C. The temperatures are not correlated with calculated jadeite or aegirine in pyroxene or with calculated
Fe3þ/Fe2þ in garnet (Table 6). Moreover, the correlation
between calculated Fe3þ/Fe2þ in garnet and pyroxene is
evidence that the assumptions used to calculate ferric iron
have validity. Garnet–chlorite temperatures were calculated by two approaches: that of Grambling (1990) and
1912
237
86
99
146
155
13
13
52
81
148
22
3
28
163
201-4E
201-4A
201-2A
201-2C
201-2B
201-2E
201-3B
201-3E
201-3G
201-3H
201-3K
201-3L
201-3N
135
G201-A
201-4D
60
G201-C
73
62
G201-H
201-4C2
64
G201-M
123
33
G201-P
201-4C1
16
G201-N
129
13
G201-F
201-4G
3
G201-G
151
245
G201-D
201-4H
Th (ppm)
Zircon
1913
135
120
35
117
63
166
98
286
282
175
239
164
201
98
126
436
238
267
111
214
98
150
226
126
83
46
87
U (ppm)
0.01557
0.01070
0.01035
0.01302
0.01013
0.01093
0.01262
0.01067
0.01104
0.01376
0.01041
0.01056
0.01300
0.02149
0.01119
0.01141
0.01147
0.00981
0.01082
0.01336
0.01087
0.01109
0.01150
0.01213
0.00906
0.01231
0.01245
0.16
0.13
0.15
0.43
0.63
0.28
1.22
0.57
0.54
0.28
0.58
2.42
0.43
0.60
0.61
0.89
0.05
0.05
0.53
0.49
2.35
0.19
0.09
0.23
1.21
U
238
Pb/
206
2.82
0.07
Th/U
1.77
3.31
1.95
2.87
1.81
2.21
1.89
1.48
1.57
1.31
1.53
1.61
1.58
1.62
1.68
1.63
1.42
1.90
1.45
1.82
1.78
1.78
1.83
2.42
2.00
1.73
2.43
%RSD
0.07986
0.06648
0.08339
0.07594
0.07967
0.07195
0.07327
0.07271
0.09029
0.07414
0.06452
0.07222
0.07451
0.08310
0.18320
0.06608
0.06688
0.07124
0.10448
0.08476
0.07160
0.06449
0.07204
0.06888
0.08390
0.11043
0.04976
U
Pb/
235
207
6.35
13.73
5.80
10.28
5.01
7.12
5.51
3.16
3.62
2.59
3.36
4.36
4.27
5.48
3.80
3.98
2.89
5.99
3.02
5.62
4.69
4.92
6.47
7.96
5.65
4.62
15.39
%RSD
0.04654
0.05323
0.04913
0.04787
0.04764
0.04800
0.04790
0.04873
0.04902
0.04689
0.04771
0.04685
0.04738
0.04635
0.06184
0.04605
0.04595
0.04681
0.05508
0.04874
0.04870
0.04619
0.04782
0.04829
0.04676
0.05146
0.03374
Pb
Pb/
206
207
Table 4: U–Pb analyses and calculated ages of zircon in garnetite GR1-201
6.57
14.05
5.98
10.59
5.16
7.33
5.70
3.22
3.71
2.60
3.44
4.46
4.35
5.72
3.93
4.30
2.92
6.15
3.10
5.74
4.83
5.04
6.73
8.18
5.82
4.72
15.50
%RSD
0.00396
0.00464
0.00328
0.00535
0.00325
0.00282
0.00305
0.00376
0.00418
0.00396
0.00253
0.00387
0.00363
0.00414
0.00890
0.00341
0.00389
0.00351
0.00527
0.00408
0.00343
0.00315
0.00347
0.00247
0.00355
0.00592
0.00200
Th
Pb/
232
208
2.27
24.14
9.15
3.55
8.92
7.09
5.90
2.39
2.39
2.02
5.53
3.62
3.03
2.66
1.57
3.52
2.83
3.13
2.47
3.92
5.25
7.94
2.88
15.38
10.70
1.69
58.00
%RSD
U
Pb/
80
79
58
78
74
71
70
86
1
2
2
1
2
1
2
1
1
1
62.9
69
1
1
1
2
1
1
1
1
1
1
1
1
1
2
2
2
2
1s
73
73.5
72
137
83
67
67.7
88
71
68
81
70
65
83
66
69
100
238
206
Ages in Ma
78
81
65
78
74
72
71
88
71
63
73
73
71
171
81
66
65
101
70
70
83
71
63
82
68
49
106
U
Pb/
235
207
5
5
9
4
7
4
5
3
2
2
2
3
3
6
4
2
3
3
4
3
4
4
3
4
5
7
5
1s
25
154
339
81
93
94
99
149
135
85
44
68
42
669
16
5
415
40
133
135
90
8
37
114
145
96
248
78
175
88
117
55
47
51
36
66
65
53
125
33
91
42
94
75
94
151
71
85
132
74
184
261
1s
125
Pb
Pb/
206
207
80
66
94
66
108
62
57
84
76
51
80
73
78
179
84
78
69
106
71
69
82
70
64
72
50
40
119
Th
Pb/
232
208
2
6
23
6
4
4
4
2
2
3
2
2
3
3
2
2
2
3
2
4
3
2
5
8
8
23
2
1s
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
JOURNAL OF PETROLOGY
ted
0.2828
ma
ntl
CH
0.2828
e1
76
Lu
UR
0.2824
(0.
7H
f
033
2)
=0
.03
84
TDM =
1.87 Ga
0.5
1.0
By before present
(b)
176Hf / 177Hf
1.5
Two or more analyses of individual zircon grains
Fig. 10. Values of 176Hf/177Hf in corresponding interior and outer
portions of individual zircon grains in garnetite GR1-201. Values for
outer portions were obtained both by laser ablation at zircon surfaces
near grain rims and by ablating through grains until garnet was
encountered. The two points representing outer portions of some
grains represent the two types of analysis.
15 My
CHUR (0.0332)
0.2828
0.2822
0.2816
2.0
DM (0.0384)
EPR
0.2824
0.2818
0.2812
0
Outer portions
Interiors
0.2820
Zircon (0.000015)
0.2816
}
SEPTEMBER 2005
0.2826
/17
0.2820
0.2832
NUMBER 9
0.2830
ple
/ 177Hf
De
0.2832
176Hf
176Hf / 177Hf
(a)
VOLUME 46
10 My
0.2824
5 My
176Lu/177Hf = 4
0.2820
Zircon (0.000015)
0.2816
0
40
80
My before present
120
Fig. 9. 176Hf/177Hf plotted against 206Pb/238U ages from the analyses
of zircons in rock GR1-201. Continuous lines and accompanying
176
Lu/177Hf values model the evolution of depleted mantle (DM) and
chondrites (CHUR), as described by Griffin et al. (2000). The bracket
labeled ‘EPR’ spans the range measured for basalts of the East Pacific
Ridge by Chauvel & Blichert-Toft (2001). Dashed lines model the
evolution of Hf in zircon with 176Lu/177Hf equal to the measured
value of 0000015 and of 176Lu/177Hf ¼ 4, consistent with the
Lu/Hf analyses of garnet in the host rock GR1-201.
that of Dickenson & Hewitt (1986) in the modified form
described by Laird (1988). All iron was considered ferrous
in both garnet and chlorite for these approaches. Both
chlorite–garnet methods yield very similar temperatures
in the range from 400 C to 500 C, and the average of the
five values is about 470 C. The contrast between the
lower garnet–chlorite and the higher garnet–pyroxene
temperatures may be due to at least two causes other
than disequilibrium—assumptions regarding ferrous
iron or faulty calibrations of thermometers for these
rocks. The garnet–pyroxene temperatures are sensitive
to assumptions made in calculating ferric and ferrous iron
in such magnesian pyroxenes, as emphasized by Proyer
et al. (2004). The garnet–chlorite thermometers were calibrated from garnet–biotite thermometry of metamorphosed pelites in which garnet and chlorite compositions
are unlike those in the garnetites. Hence, the calculated
temperatures are best regarded only as evidence that the
temperatures of garnetite formation were low, probably
in the range from 400 C to 600 C. Even if inaccurate,
the calculated values are meaningful for comparisons
of temperatures calculated by the same methods for
eclogite and pyroxenite xenoliths included in the Navajo
diatremes.
Pressures of garnetite formation are difficult to constrain. Comparisons with metamorphosed rodingites in
alpine peridotites are complicated by hydration reactions
that occurred as those rocks were exhumed. For instance,
Evans et al. (1979) suggested that metarodingites in an
Alpine peridotite had eclogite-facies mineralogies at
about 800 C and 25 GPa, and that most or all of the
amphibole and epidote in those rocks developed during a
metamorphic overprint at less than 1 GPa. The absence
of amphibole in the garnetites may be a key. Schulze et al.
(1987) and Helmstaedt & Schulze (1988) observed that
a retrograde assemblage of chlorite–garnet–omphacite
formed as a consequence of hydration of some garnet
pyroxenites, whereas pargasite–chlorite formed in others.
They concluded that the chlorite–garnet–omphacite
assemblage formed at pressures greater than about
25 GPa, above the stability of amphibole. Neither
amphibole nor epidote was identified in the garnetite
xenoliths, despite the evidence for the presence of a fluid
phase during garnet growth, and hence the chlorite–
garnet–clinopyroxene assemblage in these rocks probably
also formed near or above about 25 GPa.
Relevant xenolith assemblages in
the Navajo SUM diatremes
The SUM diatremes in the Navajo volcanic field contain
unusual xenolith types that provide context for the
interpretation that the garnetites formed in metasomatic
1914
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
Table 5: Lu–Hf analyses of zircons, model ages, and corresponding U–Pb ages
Zircon grain
176
1s
176
176
TDM (Ga)
201/3 B
0.281959
0.281852
0.000020
0.000011
0.000020
0.000012
0.001177
0.000668
1.71
1.85
0.282139
0.282186
0.282359
0.000019
0.000017
0.000015
0.000015
0.000943
0.000739
0.000721
1.48
71
0.000013
0.000012
1.42
1.19
74
0.282581
0.282259
0.000021
0.000015
0.000010
0.000010
0.000541
0.000567
0.90
1.32
0.281838
0.281915
0.000017
0.000015
0.000015
0.000016
0.000918
0.000966
1.87
1.77
0.282457
0.282738
0.000058
0.000019
0.000018
0.000018
0.001379
0.001055
1.06
0.69
0.281973
0.282413
0.281903
0.000015
0.000016
0.000011
0.000015
0.000910
0.000563
0.000960
1.70
73
0.000017
0.000011
1.12
1.79
73
0.282400
0.282196
0.000020
0.000015
0.000015
0.000016
0.000960
0.000955
1.13
1.40
0.282454
0.282434
0.000025
0.000012
0.000016
0.000012
0.000880
0.000701
1.06
1.09
0.282710
0.282070
0.282437
0.000023
0.000015
0.000013
0.000013
0.000832
0.000679
0.000704
0.72
137
0.000016
0.000015
1.57
1.09
67
0.281883
0.281886
0.282651
0.000014
0.000019
0.000028
0.000770
0.001047
0.001400
1.81
88
0.000018
0.000024
1.81
0.80
71
71
0.281829
0.281862
0.000009
0.000006
1.88
1.84
100
0.282456
0.282567
0.000017
0.000021
69
0.282851
0.281985
0.282140
0.000011
0.000021
0.282305
0.281949
0.000015
0.000017
0.282382
0.282844
0.000025
0.000011
0.000014
0.000017
0.000013
0.000836
0.000832
0.000651
1.16
0.55
0.281911
0.282622
0.000009
0.000021
0.000015
0.000023
0.000851
0.001058
1.78
0.84
0.282132
0.281868
0.281946
0.000013
0.000017
0.000015
0.000019
0.000940
0.000724
0.001174
1.49
63
0.000011
0.000011
1.83
1.73
69
201/3 E interior
201/3 E rim
201/3 G
201/3 H
201/3 K-1
201/3 K-2
201/3 N-1
201/3 N-2
201/3 N-3
201/3 S
201/4 A
201/4 A rim
201/4 C-1 interior
201/4 C-1 rim
201/4 C-2 interior
201/4 C-2 rim
201/4 C-2 rim
201/4 D
201/4 G interior
201/4 G rim
201/4 H
201 A interior
201 A rim
201 E
201 D
201 G interior
201 G rim
201 F
201 N
201 P
201 M
201 H interior
201 H rim
201 H rim
201 C
201/2 A
201/2 C
201/2 B
201/2 E
Hf/177Hf
0.000010
0.000010
Lu/177Hf
Yb/177Hf
206
Pb/238U age (Ma)
70
71
78
58
58
80
80
80
66
66
81
67
0.000012
0.000005
0.000010
0.000694
0.000210
0.000430
1.06
0.91
0.000011
0.000017
0.000014
0.000536
0.000930
0.000795
0.54
66
1.68
1.48
83
1.26
1.73
69
65
70
81
68
74
86
interior, interior of zircon grain; rim, outer portion of zircon grain.
reaction zones. These rock types include those peridotites
and garnet pyroxenites that contain chlorite and other
hydrous minerals and the eclogites. As emphasized by
Helmstaedt & Schulze (1988), the diverse lithologies are
similar to those in some metamorphosed ophiolite complexes in high-pressure orogenic belts.
Hydrous minerals in spinel peridotite xenoliths
include amphibole, chlorite, titanoclinohumite, and
antigorite: Smith (1979) concluded that all but antigorite
formed by hydration reactions in the mantle, and that at
least some of the antigorite may be of similar mantle
origin. Peridotite xenoliths also contain chlorite attributed to hydration of mantle at greater depths, below the
spinel–garnet transition (Mercier, 1976; Smith, 1995).
An unusual rock interpreted by Smith (1995) to record
reactions of garnet peridotite with water contains
1915
JOURNAL OF PETROLOGY
0.2824,
0.2828
VOLUME 46
81 Ma
NUMBER 9
SEPTEMBER 2005
68 Ma
60 µm
60 µm
0.2819
(c)
(b)
0.2819
(a)
50 µm
0.2819
60 µm
100 Ma
69 Ma
0.2825,
0.2826
(d)
(e)
100 µm
Fig. 11. Images of zircons in garnetite GR1-201 polished thin sections, together with positions of laser pits and corresponding 176Hf/177Hf ratios
and U–Pb ages. For those pits with two 176Hf/177Hf ratios, the second value represents the rim composition obtained by coring through the zircon.
(a) Cathodoluminescence (CL) image of zircon H. (b) CL image of zircon C. (c) CL image of zircon D. (d) Transmitted light image of zircon G. (e)
CL image of zircon G.
Table 6: Temperatures recorded by garnet–chlorite and garnet–pyroxene pairs and compositional parameters
T ( C)
T ( C)
% jadeite
% aegirine
Fe3þ/Fe2þ
Fe3þ/Fe2þ
T ( C)
T ( C)
CpxGar1
CpxGar2
in Cpx3
in Cpx4
in Cpx
in Gar
ChlGar5
ChlGar6
GR-41
549
492
14
6
—
—
528
461
5
2
0.05
0.02
—
N379b-GR-2
1.1
0.31
—
—
—
N379b-GR-3
690
602
6
1
—
—
618
548
12
5
0.01
0.05
—
GR1-2037
0.14
0.8
—
—
—
GR1-2037
606
531
25
6
411
631
541
4
1
0.05
0.03
402
ATG-GRP1-B
1.4
0.13
492
497
N379b-GR-4
—
—
—
—
—
—
—
—
—
—
—
497
501
GR-P1-ATG
—
—
—
—
—
—
—
—
—
—
—
431
441
N375-GR
—
—
—
—
—
—
—
—
—
—
—
492
497
Xenolith
1
Procedure
2
Procedure
3
Calculated
4
of Krogh (1988) with ferric iron calculated as discussed in text.
of Krogh Ravna (2000) with ferric iron calculated as discussed in text.
equal to total Al in a four-cation formula.
Calculated equal to (NaAlCr) in a four-cation formula.
5
Procedure of Dickenson & Hewitt (1986) in form published by Laird (1988).
6
Procedure of Grambling (1990).
7
The two clinopyroxene grains present are compositionally distinct.
centimeter-scale volumes with contrasting proportions of
chlorite, clinopyroxene, orthopyroxene, ilmenite, and
titanian chondrodite. Minerals in these volumes appear
to have formed in metasomatic reaction zones, consistent
with element transport analogous to that during rodingite
formation.
Chlorite appears together with a second generation of
garnet and clinopyroxene in some garnet pyroxenite
1916
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
xenoliths (Helmstaedt & Schulze, 1979; Benoit &
Mercier, 1986; Schulze et al., 1987; Smith, 1995).
Chlorite–garnet thermometry records temperatures in
the range from 400 C to 500 C for the pairs in apparent
textural equilibrium (Smith, 1995), just as recorded by
garnet–chlorite pairs in the garnetites (Fig. 7). Garnets
formed with chlorite in the pyroxenites have compositions near Gr28Prp36Alm36, similar to those in some
garnetites (Fig. 2), as also noted by Helmstaedt & Schulze
(1988). However, clinopyroxene in the retrograde
assemblages of the garnet pyroxenites has 37–47% Jd
component, more sodic than the range from 3% to
25% found for pyroxene in the garnetites.
Eclogite xenoliths record garnet–pyroxene temperatures between about 500 C and 700 C (e.g. Helmstaedt
& Schulze, 1988; Smith et al., 2004), similar to the values
calculated for the garnetites (Table 6). Coesite has been
identified in one eclogite, establishing a minimum pressure of about 25 GPa (Usui et al., 2003). Three eclogites
record pressures in the range 26–34 GPa by garnet–
phengite–omphacite barometry, and the calculated pressures and temperatures are consistent with the presence
of lawsonite in these rocks (Smith et al., 2004). Textures of
the eclogites indicate the presence of a water-rich fluid
during recrystallization (Smith & Zientek, 1979), and
the formation of sodic eclogite and associated jadeite
and omphacite pyroxenite has been attributed to hydrous
metasomatism (e.g. Helmstaedt & Schulze, 1988;
Wendlandt et al., 1993).
GENESIS OF THE GARNETITES
The garnetites are unlikely to be igneous in origin. First,
they record only subsolidus conditions. Chlorite, all of
which appears primary, occurs in five of the xenoliths,
and these rocks are otherwise typical of the suite. Temperatures calculated from garnet–chlorite and garnet–
clinopyroxene pairs by all methods are in the range
400–700 C, and those calculated by the more recent
methods are in the range 400–600 C (Table 6). Many
garnets contain compositional gradients that are sharp on
a scale of several micrometers, as documented by the
back-scattered electron images in Fig. 3 and by the traverses plotted in Fig. 5. Such compositional gradients
would have at least partly annealed, if cooling had been
slow enough to reset the garnet–clinopyroxene and
garnet–chlorite temperatures. Therefore, the mineralogy,
geothermometry, and textures are metamorphic. Second,
although the temperature constraints could be consistent
with low-temperature metamorphism of an igneous protolith, the xenolith compositions are unlike those of the
rare garnetites attributed to igneous processes. The igneous garnetites of mantle origin described by Exley et al.
(1983) and Zhang et al. (2004) are more magnesian than
the xenoliths. Those interpreted as deep-crustal restites
by Rivalenti et al. (1997) contain less calcic garnet. Thus
there is no evidence that the unusual bulk compositions of
the Navajo garnetites formed by an igneous process.
The bulk compositions, mineralogies, and textures of
the garnetites instead are attributed to formation in metasomatic reaction zones at contacts of mafic rock and
peridotite within the continental mantle lithosphere.
The reactions are inferred to be a consequence of fluid
flow from hydrating peridotite into the mafic rock,
a process analogous to that which forms rodingite.
Rocks that are monomineralic, or nearly so, are commonly formed in metasomatic zones (Thompson, 1959).
Coleman (1967) and many others have described rodingites that formed by metasomatism at contacts of serpentinized peridotite. The fracture-related garnet growth
establishes that a hydrous fluid was present during garnetite formation, as required for the process. Although
the garnets are unlike those of rodingites, the differences
may be due to relatively higher temperatures and pressures of garnetite formation. REE patterns of the garnets
are also consistent with formation in metasomatic reaction zones. The patterns are not like those for many
mantle garnets, in which ‘convex-up’ chondrite-normalized plots decrease smoothly from Lu to La. Rather, in
four of the five rocks, abundances of the middle REE
(MREE) are relatively high, more similar to the
‘sinusoidal’ or ‘humped’ patterns interpreted as consequences of metasomatism by Hoal et al. (1994), Roden
& Shimizu (2000), Burgess & Harte (2004), and Zhang
et al. (2004).
Garnetite formation during serpentinization is consistent with the stability of antigorite in the mantle. Aluminous antigorite is stable to temperatures above 660 C at
2 GPa (Bromiley & Pawley, 2003), hotter than almost all
temperatures calculated for the garnetites (Table 6).
Bromiley & Pawley (2003) also demonstrated that antigorite with 31 wt % Al2O3 is stabilized to higher temperatures than is pure antigorite, and antigorite in Navajo
peridotite xenoliths contains as much as 39 wt % Al2O3
and 12 wt % Cr2O3 (Smith, 1979).
Compositions of the minor minerals are consistent
with the hypothesis that peridotite was involved in the
metasomatic process. Clinopyroxene has calculated
Fe2þ/(Fe2þ þ Mg) in the range 005–008. Chlorite and
phlogopite have Fe/(Fe þ Mg) of 005–008 and 010,
respectively. These ratios are similar to those in some
peridotites. The minor minerals also have appropriate
Ni contents. Analyzed chlorite has NiO in the range
024–044 wt %, with an average of 034 wt %. For
comparison, chlorite in equilibrium with olivine in alpine
peridotites has NiO in the range 019–023 wt % in the
rocks studied by Trommsdorff & Evans (1969, 1974) and
Smith (1979).
The ranges of mineral compositions and the varieties of
compositional zoning within garnet may record sample
1917
JOURNAL OF PETROLOGY
VOLUME 46
positions within individual reaction zones and the evolution of the zones with time. MgO and Cr2O3 are zoned to
higher values at rims of some garnets (Figs 4a and 5b),
consistent with changes expected as a mafic lithology is
infiltrated by water involved in peridotite hydration.
Other features of the zonations are less easy to interpret.
For instance, CaO is relatively low in late-stage garnet in
one xenolith (Fig. 5a) and relatively high in another
(Fig. 5b). The ranges in trace element composition may
be due to the presence of gradients of chemical potential
at a variety of scales, as discussed for metarodingite
formation by Frost (1975): Zr concentrations in garnet
(Table 2) range from a low of 2 ppm in the rock with
relatively abundant zircon (GR1-201) to a high of 17 ppm
in a rock in which zircon was not found (ATG-GRP1-B).
Rodingite formation commonly is interpreted as a process that occurs in the upper crust (e.g. Evans, 1977), but
the ages deduced from U–Pb and Hf isotopic data instead
are compatible with geochronology of the Plateau mantle. U–Pb ages establish that much of zircon in garnetite
GR1-201 formed from 85 to 60 Ma, but no significant
Phanerozoic thermal event has been recognized in geochronological studies of crustal xenoliths in the Navajo
diatremes (Condie et al., 1999; Selverstone et al., 1999;
Crowley et al., 2004). Most of the age range recorded
by concordant garnetite zircon is within the 81–33 Ma
period of growth of concordant zircon in the associated
eclogite xenoliths (Usui et al., 2003; Smith et al., 2004),
and eclogite recrystallization has been attributed to mantle processes, either in the subducted Farallon slab (Usui
et al., 2003) or in the overlying mantle wedge (Wendlandt
et al., 1996; Smith et al., 2004). Garnetite genesis within
the Farallon slab is precluded by the U–Pb and Hf
isotope data that establish inheritance of Proterozoic
zircon, but the data are consistent with garnetite formation in the Proterozoic mantle wedge during Farallon
subduction.
EVOLUTION OF THE MANTLE
BELOW THE DIATREME
Age and stability of lithosphere
The inherited zircon in garnetite GR1-201 adds an
important constraint for the evolution of the lithosphere
of the Colorado Plateau. The eclogites with the most
basalt-like compositions have Nd model ages of 27 Ga
(Roden et al., 1990) and 15–18 Ga (Wendlandt et al.,
1993), but interpretation of these ages is hindered by the
REE metasomatism that affected the suite. Re-depletion
Os model ages of peridotite xenoliths from a minette plug
in the Navajo field range from 11 to 18 Ga; the range
may be due to addition or loss of Re before eruption (Lee
et al., 2001). The Hf depleted mantle model age of at least
18 Ga for garnetite zircon is a more robust indicator of
0
NUMBER 9
0
200
SEPTEMBER 2005
400
°C
600
800
1000
Crust
50
Mantle lithosphere
1
Antigorite out
2
100
40
3
20
10
150
km
0
Farallon slab
a
Fig. 12. Schematic illustration of simplified models of the lithosphere
below the Garnet Ridge diatreme. Dashed lines show calculated
geotherms. That labeled ‘0’ is a steady-state geotherm calculated to
represent the lithosphere before the beginning of subduction. The
three other geotherms are for 10, 20, and 40 Myr after the base of
the lithosphere has been truncated at 150 km depth by the Farallon
slab during flat subduction. The bold continuous line represents the
high-temperature stability limit of aluminous antigorite determined by
Bromiley & Pawley (2003). The shaded area around point 1 is appropriate for the formation of garnetite sample GR1-201, as constrained
by the concordant U–Pb zircon ages from 85 to 60 Ma. Points 2 and 3
represent eclogite xenoliths with dated zircons, as discussed in the text.
when the Plateau mantle was stabilized. That age is
matched by the oldest Hf depleted mantle model ages
of zircon cores in xenoliths from the lower crust, also
18 Ga (Crowley et al., 2004), and it is consistent with
Nd model ages of crustal xenoliths that establish the
main crust-forming event at about 185 Ga (Wendlandt
et al., 1993). The correspondence of ages for mantle and
crustal formation is evidence that the crust and uppermost mantle have been coupled together since initial
crustal formation.
Thermal histories and a tectonic model
Calculated temperature histories provide insights into the
depth of garnetite formation. A possible pre-Laramide
geotherm (Fig. 12) was calculated for a lithosphere of
200 km thickness that has a basal temperature of 1300 C
and a surface heat flow of 56 mW/m2; the heat flow is the
average of the two values closest to the diatreme in the
compilation of Minier & Reiter (1991). Geotherms also
are plotted for a model of the lithosphere during flat
subduction of the Farallon slab, 700 km from the trench
with a relative plate convergence of 10 cm/year. The
geotherms for the period during Farallon subduction
were calculated using a standard finite difference conductive thermal model similar to but simpler than that
used by Spencer (1996). He calculated geotherms for
models in which the Farallon slab sheared off and
replaced the lower part of the Plateau lithosphere: the
most significant difference in the model used here is that
the geotherm in the Farallon plate at the trench was not
adjusted for the age of the subducting slab. The interplate
1918
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
contact was assumed to be at 150 km: this thickness for
the remnant lithosphere is consistent with the minimum
value of 120 km based on analyses of Os in xenoliths
erupted at about 25 Ma (Lee et al., 2001) and with the
value of 120–150 km based on interpretations of data
from the LA RISTRA seismic line (West et al., 2004). The
LA RISTRA array passes about 5 km from the diatreme.
The crustal thickness of about 47 km is constrained by
receiver function analysis of data from that same seismic
array (Wilson et al., 2003). The model used for geotherm
calculation is simplified and the necessary assumptions
are not well constrained, but one conclusion is relatively
robust: before and near the beginning of low-angle subduction, and if temperatures were determined only by
conductive heat flow, then antigorite was stable only in
the upper few tens of kilometers of the continental mantle
lithosphere.
The oldest nearly concordant U–Pb ages of zircon
garnetite xenolith GR1-201 cluster near 85 Ma (Table 4,
Fig. 8). In contrast, flat subduction has been inferred to
have occurred during the Laramide orogeny, from about
80 to 40 Ma (Coney & Reynolds, 1977; Spencer, 1996).
Geological relationships on the Colorado Plateau also are
consistent with the beginning of the Laramide orogeny at
about 80 Ma. Intrusions in the Carrizo Mountains, about
50 km east of the Garnet Ridge diatreme, were emplaced
from 74 to 71 Ma and have been ascribed to Laramide
processes (Semken & McIntosh, 1997). In northern New
Mexico east of the Carrizo Mountains, stratigraphic and
structural effects of the Laramide orogeny began in the
period from 80 to 75 Ma (Cather, 2004). If the model
lithosphere and geotherms in Fig. 12 reproduce conditions before and during the Laramide orogeny, then at
times before and within about 15 Myr after the beginning
of flat subduction, mantle serpentinization could occur
only at depths shallower than about 85 km. If so, and if
the garnetites are rodingite analogues, then they must
have formed in the uppermost mantle.
The actual history of the continental lithosphere below
the diatremes during Farallon subduction may have been
much more complex than that assumed for construction
of Fig. 12. Complexities in the mantle today are illustrated
in a model derived from compressional and shear seismic
phases recorded by the LA RISTRA seismic array (Gao
et al., 2004). Their tomographic reconstructions have
steeply dipping boundaries in seismic velocity that extend
beneath the Navajo volcanic field from 50 km to 200 km
depth, and they suggest that these boundaries are determined by temperature differences related to Proterozoic
structures. The presence of a mantle suture below the
diatreme is consistent with the suggestion by Selverstone
et al. (1999) that the eclogite xenoliths are from a Proterozoic subduction zone. Fluid infiltration and serpentinization above subducting slabs may be controlled by
structural features (Hyndman & Peacock, 2003). The rise
of buoyant serpentinite masses, such as those discussed
by Guillot et al. (2001) and Ueda et al. (2004), could also
be controlled by existing structures. The conductive geotherms plotted in Fig. 12 provide only a starting point to
consider depths of origin of the garnetite xenoliths, as the
temperatures in and near the probable suture zone may
have been influenced by the flow of serpentinite and of
hydrous fluids.
Relevance to the genesis of associated
eclogites
Formation of the garnetites in metasomatic reaction
zones demands the presence of one or more other rock
types to react with peridotite, and the eclogite xenoliths
represent possible reactants. Both garnetite GR1-201 and
some of the eclogites inherited zircon from Proterozoic
protoliths and also contain zircon that grew during the
interval from 85 to 33 Ma (Usui et al., 2003; Smith et al.,
2004). Both rock types retain evidence of protoliths with
low-pressure histories. The eclogites probably represent
subducted oceanic crust, because of their oxygen isotope
compositions (Smith et al., 2004), and because a negative
Eu anomaly in one eclogite has been attributed to plagioclase fractionation (Roden et al., 1990). The chondritenormalized REE pattern (Fig. 6) for one garnetite has a
positive Eu anomaly that also may be inherited from
subducted oceanic crust. Although the eclogites contain
more jadeite-rich clinopyroxene than do the garnetites,
the contrast is like that between pyroxenes in genetically
related eclogites and metarodingites described by Evans
et al. (1979, 1981).
The possibility that garnetite and eclogite xenoliths
were from the same mantle source region can be tested
by comparisons of pressures and temperatures recorded
by two Navajo eclogites that contain dated zircons. Pressures and temperatures calculated by Smith et al. (2004)
for these two rocks yield depths greater than those outlined for garnetite formation in Fig. 12. One of the xenoliths contains a fraction of concordant zircon that grew
at about 39 Ma (Smith et al., 2004), near the end of the
Laramide orogeny, and about 10 Myr before diatreme
eruption. That eclogite plots at a temperature slightly
higher than the boundary determined for antigorite stability by Bromiley & Pawley (2003) (Fig. 12, point 3). The
other eclogite contains zircon that yielded concordant
ages from 81 to 47 Ma (Usui et al., 2003), and it plots at
conditions too cool for any of the calculated geotherms,
even after 40 Myr of low-angle subduction (Fig. 12,
point 2). The thermobarometry may be inaccurate, and
the model may not reproduce the temperature evolution
of the mantle. If so, the calculations and the model do
not preclude a common source region for the xenoliths,
despite these discrepancies. Regardless, the similar zircon
chronologies and cool garnet–pyroxene temperatures are
1919
JOURNAL OF PETROLOGY
VOLUME 46
evidence that garnetite and eclogite record a common
process related to rock–fluid interactions.
Hydration of the mantle lithosphere
Formation of the garnetites as rodingite analogues in
Proterozoic mantle demands the presence of water, and
at least two hydration episodes are recorded by the xenoliths and xenocrysts. Pyrope grains are scattered throughout the SUM host rock at Garnet Ridge, and some of
these garnets contain inclusions of hydrous minerals such
as chlorite, titanoclinohumite, and carmichaelite as well
as of olivine and pyroxene (McGetchin & Silver, 1970;
Hunter & Smith, 1981; Wang et al., 1999). The pyrope
grains with included chlorite have complex REE patterns
similar to some of those plotted in Fig. 6 (Roden &
Shimizu, 2000). The included hydrous minerals appear
to have been trapped in pyrope during prograde garnet
growth. In contrast, chlorite, amphibole, antigorite, and
other hydrous minerals also formed during retrograde
hydration of spinel and garnet peridotite and pyroxenite
(Mercier, 1976; Helmstaedt & Schulze, 1979; Smith,
1979, 1995). The retrograde hydration is more plausibly
related to water introduced into the mantle wedge
during low-angle Farallon subduction accompanying
the Laramide orogeny. In is unclear which hydration
event, if either, is related to garnetite formation and the
growth of concordant zircon in garnetite and eclogite at
about 80 Ma.
The volume of mantle hydrated during the Laramide
orogeny cannot be constrained from these xenoliths. The
hydrated sources of the SUM diatremes must have been
deeper than those of the xenoliths, and so below about
110 km, if the calculated depths and temperatures plotted
in Fig. 12 are accurate. Smith et al. (2004) suggested that
the hydration might have been restricted to long-lived
tectonic boundaries within the mantle, like those discussed by Selverstone et al. (1999) and Gao et al. (2004).
In contrast, Humphreys et al. (2003) suggested that
dehydration of the Farallon slab resulted in such extensive hydration of the overlying mantle that it caused
regional uplift of the western USA. Regardless of the
extent of hydration, the garnetites provide evidence of
at least local hydrous metasomatism in the continental
mantle near the beginning of the Laramide orogeny.
If the hydration was due to water released from the
Farallon slab or mobilized by associated magmatism,
then the oldest concordant U–Pb ages of about 85 Ma
may be useful in constraining the history of that slab.
NUMBER 9
SEPTEMBER 2005
and Smith et al. (2004) scatter over a period of at least
30 Myr. However, the presence of inherited components
of clearly older ages makes the U–Pb data ambiguous in
terms of the timing of the metasomatic events that produced the eclogites and garnetites, because any given
grain (or a given ablated volume) might contain a mixture
of inherited and metamorphic zircon, giving mixed ages.
U–Pb data alone also cannot distinguish between the
growth of new zircon and complete loss of Pb from
older zircons. Zheng et al. (2004a, 2004b) have demonstrated how Hf isotope data can resolve these ambiguities.
The Hf isotope analyses for GR1-201 provide a much
better estimate of the age of the protolith than was available from U–Pb data, and they provide a rough estimate
of the duration of the metasomatic event(s). The wide
spread in 176Hf/177Hf is consistent with the rapid evolution of radiogenic Hf in the garnetite matrix (Fig. 9).
Scherer et al. (1997) suggested that the presence of zircon
would buffer the Lu/Hf of a garnet granulite to low
values. This clearly was not the case in garnetite GR1201, as Hf has been sequestered into zircon and Lu into
garnet, leaving the garnet matrix with an extremely high
Lu/Hf and the zircon with an extremely low Lu/Hf. The
migration of radiogenic Hf from the garnet into existing
zircons (producing the observed Hf-isotope zoning), and
its incorporation into several generations of newly grown
zircon, would be enhanced by the periodic movement of
fluids through the matrix that is suggested by the majorelement zoning within garnet (Figs 3 and 5).
The 15 Myr required for the evolution of the observed
spread in Hf isotope ratios (Fig. 9) may be a minimum
estimate for the time over which episodic metasomatism
and zircon growth took place within the garnetite. Some
of the scatter about the growth curve (Fig. 9) may reflect
the mixing of domains with different ages and different
Hf-isotope compositions, at the scale of the laser-ablation
analyses. Assuming that the most radiogenic Hf-isotope
compositions have not been measured because of mixing
of domains during ablation, the metasomatic episodes
may have spread over a longer time, such as the 20–
25 Myr reflected by the main spread of zircon ages
(Fig. 8). However, it seems clear that the metasomatism
that produced the garnetites occurred over a short period
(15–25 Myr) near the end of the Cretaceous and the
beginning of the Laramide orogeny, and that garnetite
formation overlapped the period of formation of some
zircons in the associated eclogites.
Implications for processes in the
mantle wedge
DISCUSSION
Timing and duration of metasomatism
The concordant zircon U–Pb ages in sample GR1-201
and in the eclogite samples studied by Usui et al. (2003)
The garnetite provides unusual evidence for rock–water
interactions in continental mantle. Except for examples
from the Colorado Plateau, most xenolith evidence for
such hydration is restricted to the presence of amphibole
1920
SMITH AND GRIFFIN
GARNETITE XENOLITHS, COLORADO PLATEAU
and phlogopite, as summarized by Luth (2004). Even
though serpentinization of continental mantle may be a
common and important process in the forearc wedge
(Hyndman & Peacock, 2003), most evidence for it is
based on seismic data. The hypothesis that low-angle
subduction has caused extensive and widespread hydration of continental lithosphere (Humphreys et al., 2003) is
difficult to test. If the garnetites are rodingite analogues,
then they document serpentinization above the Farallon
slab and at about 700 km from the trench, consistent with
that hypothesis. The zircon U–Pb ages establish that
garnetite formation began 5–10 Myr before Laramide
tectonism, however, and the timing may indicate that
hydration and low-angle subduction can precede related
tectonism and magmatism by that time interval. Calculated pressures and temperatures together with zircon
U–Pb ages do not fit simple thermal histories of conductive heat transfer in the mantle above the Farallon slab
(Fig. 12). Coupled geochronological and petrological
analyses of additional xenoliths from the Navajo diatremes may document how fluid flow and serpentine
diapirism can affect the mantle wedge.
The garnetite xenoliths also provide evidence of mobility of elements that are important in deciphering mechanisms of arc volcanism. Some of these elements, such as
Ti, Zr, and Hf, have been termed ‘conservative’ and
immobile in aqueous fluids (Pearce & Peate, 1995), particularly in the presence of rutile (Brenan et al., 1994).
Evans et al. (1981) observed that Ti, Zr, and Hf appeared
to have been immobile in formation of metarodingites,
but such immobility may not be the rule at either crust
or mantle pressures. Zr and Ti can be mobile during
rodingite formation at low pressures (Dubinska et al.,
2004) and at relatively high pressures during formation
of eclogite veins containing zircon and rutile (Philippot &
Selverstone, 1991; Rubatto & Hermann, 2003).
Woodhead et al. (2001) found evidence for nonconservative behavior of Hf in slab–wedge interactions, perhaps
because of fluid transport of Hf, a suggestion they viewed
as radical. The Hf isotope abundances and the U–Pb
ages testify to the mobility of Zr and Hf in formation of
garnetite GR1-201.
ACKNOWLEDGEMENTS
Some of the garnetite samples were collected by
W. C. Hunter and by A. T. Gavasci. LAM–ICPMS
analyses at The University of Texas at Austin were
acquired with the assistance of J. Lansdown. Norm
Pearson and Suzie Elhlou are thanked for their cheerful
and patient assistance with the U–Pb and Lu–Hf analyses. For many years, D. J. Schulze has shared information about Navajo garnetite and eclogite xenoliths, and
insights into their petrogenesis, much to the benefit of
this contribution. M. F. Roden made valuable comments
during manuscript preparation and on a preliminary
draft. This paper was much improved as a result of
constructive reviews by J. Selverstone, J. G. Liou and
C. Shaw, and by editor B. R. Frost. This is publication
383 from the ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents
(www.es.mq.edu.au/GEMOC). Support for research was
provided by the Geology Foundation of The University
of Texas at Austin.
SUPPLEMENTARY DATA
Supplementary data for this paper are available at Journal
of Petrology online.
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