Garnetite Xenoliths and Mantle–WaterInteractions Below the Colorado Plateau,Southwestern United States

DOUGLAS SMITH1* AND WILLIAM L. GRIFFIN2,3

1DEPARTMENT 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

2GEMOC KEY CENTRE, DEPARTMENT OF EARTH AND PLANETARY SCIENCES, MACQUARIE UNIVERSITY, SYDNEY,

NSW 2109, AUSTRALIA

3CSIRO EXPLORATION AND MINING, NORTH RYDE, NSW 2113, AUSTRALIA

RECEIVED JULY 7, 2004; ACCEPTED MARCH 21, 2005ADVANCE 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 composi-

tions. Most U–Pb analyses plot on or near concordia in the range

60–85Ma but a few are discordant. The range in 176Hf/177Hf is

about 0�2818–0�2828, with grains zoned to more radiogenic Hf

from interiors to rims. The garnetite protolith contained zircons at

least 1�8 Ga in age, and garnet and additional zircon crystallized

episodically during the interval 85–60Ma. 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, begin-

ning 5–10Myr before tectonism related to low-angle subduction.

KEY WORDS: garnetite; Lu–Hf, mantle; rodingite; metasomatism

INTRODUCTION

Garnetite xenoliths from a diatreme cluster in the Navajovolcanic field of the Colorado Plateau have been studied

to investigate implications for serpentinization andmetasomatism of continental mantle. Helmstaedt &Schulze (1988) suggested that these xenoliths formed byreactions involving peridotite hydration. If so, they maybe samples of metasomatic reaction zones analogous tothose that form at low temperatures and pressures duringserpentinization (Coleman, 1967); garnet-rich parts ofsuch zones are called rodingites. If the garnetites areanalogues of rodingites, then they may provide evidenceof mantle processes associated with subduction of theFarallon plate, by analogy with origins proposed foreclogite xenoliths from the same diatremes. The localityis about 700 km from the Farallon subduction zoneaccording to the reconstruction of Severinghaus &Atwater (1990); however, this is more than three timesthe maximum distance from the trench that Hyndman &Peacock (2003) considered a typical limit for subduction-induced serpentinization. Evidence that the garnetites arerodingite analogues would be consistent with the hypo-thesis of Humphreys et al. (2003) that low-angle subduc-tion hydrated the mantle wedge below a broad region ofwestern North America.The garnetites also have been studied to clarify the

genesis of the associated eclogite xenoliths. The eclogiteshave been interpreted either as fragments of the Farallonslab itself (Helmstaedt & Doig, 1975; Usui et al., 2003)or as products of water–rock reactions in the mantlewedge (Smith et al., 2004). These contrasting hypotheses

*Corresponding author. Telephone: either 512-471-4261 or 970-259-

0558. E-mail: doug@geo.utexas.edu. (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

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 PAGES 1901–1924 2005 doi:10.1093/petrology/egi042

imply very different effects of low-angle subduction onthe mantle roots of continents. The garnetite xenolithsare hosted only by those diatremes that host the eclogites,and Switzer (1975) suggested that these rock types weregenetically related. If so, the garnetites yield insights intohow the eclogites formed.

OCCURRENCE

All garnetite xenoliths studied here are from the GarnetRidge diatreme cluster on the Comb Ridge monocline innortheastern Arizona (Fig. 1). The host rocks in thediatremes are serpentinized ultramafic microbreccia(SUM; Roden, 1981). These SUM host rocks appear tohave been emplaced as gas–solid mixes (McGetchin &Silver, 1972), and textural and chemical evidence sup-ports the conclusion that a melt phase was never partof the mixture (Roden, 1981). Rather, the gas–solideruptions have been interpreted as products of hydratedmantle disaggregated during heating by intruded magma(Smith & Levy, 1976). The Garnet Ridge diatremes wereemplaced at 30Ma (Smith et al., 2004).Garnetite occurs in the diatremes together with an

extraordinary variety of xenoliths of sedimentary, meta-morphic, and igneous rocks: gabbro, granite, rhyolite,granulite, amphibolite, peridotite, omphacite pyroxenite,and eclogite are among the rock types present. Eclogitemakes up less than 0�05 wt % of the population of igneousplus metamorphic rocks (Hunter, 1979). The abundanceof garnetite xenoliths was not measured, but they are lesscommon than those of eclogite, and they probably makeup less than 0�01 wt % of that population. Garnetite alsooccurs in the two other major diatremes on the CombRidgemonocline,Mule Ear andMoses Rock (Helmstaedt& Schulze, 1988; McGetchin & Silver, 1970).

DESCRIPTIONS AND

PETROGRAPHY

The largest of the 14 garnetite xenoliths studied in thinsection had a maximum diameter of 8 cm, slightly smal-ler than the 11 cm dimension of the largest garnetitexenolith described from the province by Switzer (1975).Most were less than 4 cm in maximum diameter. Typicalspecimens are nearly equant and have smooth surfaces,some of which appear polished. These smooth surfacesare attributed to abrasion and impact during emplace-ment of the gas–solid mix that formed the diatreme fill, asdescribed by McGetchin & Silver (1970, 1972). Irregularfractures, 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. Slightvariations of garnet color from clear to very paleorange are visible in sections of some rocks. These color

variations were seen only in polished thin sections withthicknesses in the range from 100 to 300mm. In somerocks, parts of garnets are turbid with unoriented inclu-sions a few micrometers in average diameter, many ofwhich appear to be of fluid. The grain boundaries ofgarnets were inferred primarily from the color variationsand turbidity and from cracks. Where estimates werepossible, garnet grains appeared anhedral and approxim-ately equant, typically with maximum diameters of a fewmillimeters.A diverse suite of minor and trace minerals is present;

multiple polished thin sections were prepared of mostsamples so as to identify as many as possible. In approx-imate order of decreasing abundance, these minerals are:rutile, ilmenite, chlorite, clinopyroxene, zircon, pyrite,phlogopite, and apatite. Rutile and ilmenite are the onlyminerals other than garnet present at a modal abundanceof 1% or more; the others are present only as traceconstituents. Grains of the two oxides typically are anhed-ral, with maximum dimensions of several millimeters orless. Rutile forms an estimated 5% of one sample, but itmakes up less than 2% of more typical samples and isabsent in some. The maximum abundance of ilmenitewas estimated as about 2%; most samples contain �1%,and it was not observed in some rocks. Chlorite wasidentified in six of the 14 samples, as was clinopyroxene.Apatite, phlogopite, and pyrite were identified in twoto four rocks. Clinopyroxene, chlorite, and apatite areslightly more common at the edges of several of thexenoliths in which they occur. In rocks that contain

500 km

109°

03' W

37°00' N

COUT

AZ NM

ColoradoPlateau

Garnet Ridge

Navajovolcanicfield

Fig. 1. Location of the study area within the Colorado Plateau. ~,location of the Garnet Ridge and other major diatremes of serpentin-ized ultramafic microbreccia (SUM) in the Navajo volcanic field. Theboundary of the Navajo field is determined by the distribution ofminette and related rocks.

1902

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

both ilmenite and rutile, they occur as discrete grainsand as rims around one another. Chlorite is intergrownwith rutile in some occurrences and with ilmenite in someothers. No other systematic mineral associations werenoticed.

Zircon was identified in five xenoliths, but only in oneof these rocks (GR1-201) were more than a few grainsobserved. Although zircon is present only in trace pro-portion, it is the most common silicate other than garnetin that rock. Zircon grains in GR1-201 were chosenfor in situ U–Pb and Lu–Hf analysis. Most grains areanhedral to subhedral, but rare grains are euhedral;maximum diameters range from several micrometers toabout 300 mm.

MINERAL CHEMISTRY

Procedures for elemental analysis

Minerals in 11 xenoliths were analyzed by wavelength-dispersive spectrometry (WDS) on JEOL 7300 and 8200electron microprobes at The University of Texas atAustin. For most analyses, the accelerating voltage was15 keV and the beam current was about 40 nA. Countingat peak and background wavelengths was terminated at40 s for each, unless a standard deviation <0�3% wasachieved first, based on counting statistics. In a few ana-lytical sessions, trace elements were analyzed with a20 keV accelerating voltage, higher beam current, andlonger times for data acquisition. Data were correctedwith a JEOL ZAF procedure. Representative electronmicroprobe (EMP) analyses of minerals in six of the xeno-liths are given in Table 1, and representative analyses forall of the xenoliths are given in the Electronic Appendix(available at http://www.petrology.oupjournals.org).

Trace elements in minerals of five rocks were analyzedby laser ablation microprobe–inductively coupled plasmamass spectroscopy (LAM–ICPMS) at The University ofTexas. After photography and characterization by EMPanalysis, minerals in polished thin sections were ablatedwith a New Wave LUV213 laser. The Nd:YAG sourcedelivers a focused 213 nm beam; resulting ablation pitshave diameters from 50 to 100mm. The ablated materialwas transported with a helium carrier gas into a Micro-mass Platform quadrupole inductively coupled plasmamass spectrometer. Signal intensities were tabulated foreach mass sweep in most of the analyses, and anomalousdata were excluded. Concentrations in garnet andclinopyroxene were calculated with 44Ca as an internalstandard in two of the three datasets and with 43Ca in theother: reanalysis verified consistency of the two choices.Two semiquantitative analyses of rutile were calculatedwith 93Nb as an internal standard. Four NIST SRMglasses (610, 612, 614, and 616) and USGS glass BCR-2G were analyzed during each session. For almost all

analyzed masses, calibration curves were fitted either toSRM 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, exceptin rare cases when it was included as a primary standardto avoid extrapolations to much higher concentrationsthan those in NIST SRM 610. Consistency of some of thegarnet and pyroxene analyses was tested and verifiedby analysis of two masses for each of five REE. Accuracywas verified by the analyses of the secondary standard,BCR-2G. Backgrounds were obtained by measurementswithout ablation of samples. Detection limits were notcalculated from these backgrounds, because of possibledifferences between signal levels with and without sampleablation. Instead, results are not reported below deter-mination limits based on average concentrations withintwo standard deviations of zero. Representative deter-mination limits were about 55 ppb for La in garnet andnear and below 12 ppb for Tm, Yb, and Lu in pyroxene.

Garnet

The most common garnet compositions have roughlyequal 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 followingaverage contents: TiO2, 0�07 wt %; Cr2O3, 0�05 wt %;MnO, 0�34 wt %; Na2O, 0�03 wt %. No clear correla-tions were observed between major (Ca, Fe, Mg) andminor elements (Na, Ti, Cr, and Mn) in the entire data-base, although Fe and Mn are well correlated in thegarnet of some rocks. The most magnesian garnets(50% pyrope end-member) are found in xenolith N379b-GR-4, and they have more Cr2O3 (0�13–0�17 wt %) thangarnet in other samples.

Many garnets are complexly zoned, as illustrated inback-scattered electron (BSE) images in Fig. 3. Two typesof zoning were distinguished. First, gradients from graininteriors to rims were observed (Fig. 3a and b). Second,compositional gradients also are present in patternsthat record both growth of garnet within fractures andreaction of the fractured grains (Fig. 3c). The more com-plex patterns formed by the interplay of both processes(Fig. 3d and e).

No trends of compositional zoning are consistent forthe full dataset, although trends in individual garnets arewell defined (Figs 4 and 5). BSE images document com-positional changes at rims adjacent to ilmenite in rockN375-GR (Fig. 3a and b): one such rim, about 400 mmwide, is Mg-rich and Ti-poor relative to the interior(Fig. 4a). Annular orange zones are present in garnets ofrock GR1-202, and a traverse across such a zone and an

1903

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

Table 1: Representative electron microprobe analyses of minerals

Sample: GR1-203

Mineral: Gar1 Chl1 Chl Gar2 Cpx2 Gar3 Cpx3 Rut Ilm

SiO2 40.2 30.8 31.2 40.2 55.9 39.6 55.4 0.06 0.06

TiO2 0.05 0.01 0.07 0.10 0.07 0.10 0.04 96.8 51.5

Al2O3 22.0 19.3 19.0 22.1 5.86 21.8 2.78 0.22 0.06

Cr2O3 n.a. n.a. 0.02 n.a. 0.01 0.04 0.03 0.11 0.03

FeOt 16.6 4.86 4.87 16.7 3.41 18.0 3.70 2.58 43.3

MnO 0.29 0.00 0.02 0.24 0.01 0.29 0.02 0.00 0.1

MgO 7.62 31.1 31.6 8.56 12.5 8.48 14.3 0.02 4.62

CaO 14.0 0.03 0.01 12.8 18.2 11.6 21.5 0.01 0.00

Na2O 0.00 0.01 0.01 0.03 4.41 0.02 2.42 0.00 0.00

NiO n.a. n.a. 0.41 n.a. n.a. 0.02 0.07 0.00 0.20

Total 100.8 86.1 87.3 100.7 100.4 100.1 100.2 99.8 99.8

Sample: N379b-GR-4 GR1-201

Mineral: Gar1 Chl1 Ilm Rut4 Rut5 Gar Gar Rut

SiO2 41.3 32.0 0.02 n.a. n.a. 39.4 39.2 n.a.

TiO2 0.05 0.00 53.8 93.4 94.9 0.07 0.02 96.5

Al2O3 23.2 18.9 0.05 0.18 0.17 21.8 22.1 0.18

Cr2O3 0.16 0.10 0.06 0.63 0.53 0.09 0.02 0.24

FeOt 14.4 3.48 39.3 2.00 1.82 17.7 23.1 1.82

MnO 0.35 0.02 0.14 n.a. n.a. 0.49 1.00 n.a.

MgO 14.0 32.6 7.10 n.a. n.a. 7.91 8.04 n.a.

CaO 8.11 0.02 0.00 n.a. n.a. 11.9 6.90 n.a.

Na2O 0.00 0.02 0.01 n.a. n.a. 0.03 0.04 n.a.

NiO 0.02 0.37 0.21 n.a. n.a. 0.00 0.00 n.a.

Nb2O5 n.a. n.a. n.a. 1.20 0.62 n.a. n.a. 0.04

Total 101.8 87.5 100.7 97.4 98.1 99.4 100.5 98.8

Sample: ATG-GRP1-B N379b-GR-1 GR1-202

Mineral: Gar1 Chl1 Gar2 Cpx2 Ilm Rut Gar6 Gar7 Gar

SiO2 40.5 32.6 40.0 54.2 0.00 0.02 39.2 39.8 39.8

TiO2 0.08 0.00 0.05 0.00 54.1 97.2 0.09 0.07 0.10

Al2O3 22.8 18.9 22.9 0.81 0.07 0.17 21.5 22.5 22.3

Cr2O3 0.01 0.08 0.05 0.02 0.08 0.29 0.02 0.07 0.02

FeOt 15.4 3.51 15.0 2.08 40.6 2.48 18.8 12.9 16.0

MnO 0.28 0.01 0.31 0.02 0.09 0.02 0.25 0.14 0.24

MgO 12.5 31.9 12.3 17.1 6.58 0.01 6.46 8.33 8.88

CaO 9.83 0.02 10.3 24.7 0.02 0.04 13.6 16.4 12.9

Na2O 0.02 0.01 0.03 0.60 0.00 0.01 0.05 0.01 0.02

NiO 0.00 0.28 0.00 0.04 0.22 0.00 n.a. n.a. n.a.

Total 101.4 87.3 100.9 99.6 101.7 100.2 100.0 100.1 100.2

FeOt, total Fe as FeO; n.a., not analyzed.1Pairs used for garnet�chlorite temperature calculations.2,3Pairs used for garnet�pyroxene temperature calculations.4,5Analyses of rutile with highest and lowest Nb in rock.6,7Interior and rim, respectively.

1904

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

apparent grain boundary is plotted in Fig. 4b; Ti is highat the boundary relative to the grain interior, unlike theexample plotted in Fig. 4a. The gradients in the garnet ofFig. 5 are due to both ‘normal’ and fracture-relatedgrowth. In a clear case of fracture-related zonation inN379b-GR-2, the fracture fill is less calcic and less titani-ferous than the surrounding garnet (Fig. 5a). Garnet inN379b-GR-1 has complex zoning that includes bothinterior-to-rim gradients and gradients along sets ofsmall intersecting fractures (Fig. 5b).Trace elements determined by LAM–ICPMS in garnet

include Ni, Y, Zr, Yb, and Hf and all the rare earthelements (REE) except La (Table 2). Measurements ofLa, Nb, Ta, and U yielded concentrations below deter-mination limits. Average Ni concentrations fall in therange 8–18 ppm. Of garnet in the five xenoliths, that inzircon-bearing GR1-201 is highest in Y (67 ppm), Lu(1�3 ppm) and the other heavy REE (HREE), and lowestin Zr (2 ppm) and Hf (0�04 ppm). Chondrite-normalizedprofiles of REE abundance in garnet in each of the fiverocks are relatively similar for the range Ce to Sm(Fig. 6a). In three of the rocks, however, normalizedREE abundances decrease with increasing atomic num-ber in the range Sm to Lu, forming ‘humped’ patterns. Inone of these three rocks, there is a distinct positive Euanomaly, and in the other two, poorly defined negativeanomalies. The patterns of average REE abundance foreach 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 inthis rock form a clustered group, whereas the sixth ismuch lower in HREE (Fig. 6b). The volume of thisanomalous analysis is distinguished by lower averageatomic weight and so is relatively dark on the BSEimage (Fig. 3d); the volume also has relatively lower Feand higher Mg, Ca, and Cr (Fig. 5b). Although nocomparable inhomogeneity in REE was documented ingarnet of the other four samples, analytically significantdifferences in concentration are present (Fig. 6b and c).

Rutile, ilmenite, clinopyroxene, chlorite,and phlogopite

Compositions of rutile and ilmenite are similar in all butxenolith N379b-GR-4 (Table 1 and the ElectronicAppendix). Ilmenite, analyzed by EMP in five xenoliths,contains 0�09–0�27 wt % NiO and 4�0–7�1 wt % MgOand trace Cr, Al, and Mn; that in sample N379b-GR-4has the highest Mg, Ni, and Cr. EMP analyses of rutile inseven xenoliths establish the typical range of 0�1–0�3wt%Cr2O3, but rutile in N379b-GR-4 contains 0�5–0�6 wt %Cr2O3. In two of the three rocks in which rutile wasanalyzed for Nb2O5, concentrations are near 0�04 wt %,but they are much higher, 0�62–1�20 wt % Nb2O5, inN379b-GR-4. Concentrations of V and Zr in rutile,determined by semiquantitative LAM–ICPMS analysis,

50

040

30

20

10

0

10 20 30 40 50

60

0

10

20

30

405060 Mg

Ca

Fe

50

040

30

20

10

0

10 20 30 40 50

60

0

10

20

30

405060

(b) Rock (# analyses)GR1-201 (29)N379b-GR-1 (51)N379b-GR-2 (39)

(a) Rock (# analyses)GR1-202 (55)GR1-203 (8)GR-P1-ATG (7)GR-41 (2)

N375-GR (14)N379b-GR-3 (4)N379b-GR-4 (5)ATG-GRP1-B (6)

Fig. 2. Fe–Ca–Mg proportions in garnets of 11 garnetite xenoliths. (a) Analyses of garnet in eight xenoliths. Garnets in these samples eitherappeared relatively homogeneous for these elements, or too few analyses were made to determine if heterogeneities are present. (b) Analyses ofgarnet in three xenoliths in which heterogeneity for these elements is well documented.

1905

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

are about 4500 and 170 ppm, respectively, in xenolithN379b-GR-4 and about 3000 and 60 ppm in the relat-ively zircon-rich rock, GR1-201.Clinopyroxene was analyzed in five xenoliths (Tables 1

and 2, and Electronic Appendix). In three of the five, onlyone small grain was found, and in the fourth, only twograins. In three xenoliths, the Jd component of the pyr-oxene is in the range 3–6%, and in one it is 14%; in thexenolith with two identified grains, one has 12% Jd andthe other 25%. Aegirine, calculated as (Na – Al – Cr) on a

four-cation basis, ranges from 0�5% to 6%. The domin-ant end-member is diopside, and calculated Fe2þ/(Fe2þþMg) of these pyroxenes ranges from 0�05 to 0�08. A rela-tively large grain of clinopyroxene, about 3mm in max-imum diameter, was found in one rock (ATG-GRP1-B),together with four much smaller grains. The chondrite-normalized pattern of REE abundances in the large grainhas 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.

100 µm

(e)

(d)

1000 µm

100 µm

(c)

100 µm

(b)

Ilmen

ite

400 µm

(a)

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 atgrain boundary with ilmenite in rock N375-GR. Diamonds mark positions of the analyses plotted in Fig. 4a. The outlined rectangle is shown withenhanced contrast in (b). (b) Contrast-enhanced portion of field of view in (a), showing irregular compositional boundaries. (c) Apparent fracture-related compositional zoning in rock N379b-GR-2; the path of the compositional traverse (Fig. 5a) is visible because of oil condensation under theelectron beam. (d) Complex zonation in rock N379b-GR-1. Garnet is the only mineral in this image. Diamonds show positions of analyses plottedin 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 EMPtraverse, within the darker (lower average atomic weight) garnet in this BSE image. (e) Fracture-related and oscillatory zoning in an enlargementof the rectangle outlined in (d).

1906

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

Chlorite, analyzed by EMP in five xenoliths, isclinochlore, with Fe/(Fe þ Mg) in the range 0�05–0�08(Table 1 and Electronic Appendix). NiO contents rangefrom 0�24 to 0�44 wt % and average 0�34 wt %. Fe/Mg ofchlorite inclusions varies systematically with that of thesurrounding garnet (Fig. 7). Phlogopite was identified intwo rocks and analyzed in one. The phlogopite hasFe/(Fe þ Mg) ¼ 0�10 and 0�22 wt % NiO.

ZIRCON GEOCHRONOLOGY

Analytical techniques

Zircons in four polished sections of garnetite GR1-201were analyzed at GEMOC in Macquarie University.U–Pb ages were determined by LAM–ICPMS asdescribed by Belousova et al. (2001) and Jackson et al.(2004). The instrumentation consisted of a modifiedMerchantek/New Wave LUV 213 nm Nd:YAG laserattached to an Agilent 4500 s ICPMS system. Typicalablation pits are 30mm in diameter, and all ablationswere performed in a He carrier gas. Mass bias andinstrumental drift were corrected using a very homo-geneous external standard, the GEMOC GJ-1 zircon(608Ma). Data were reduced using the in-house GLIT-TER software (www.es.mq.edu.au/GEMOC), which

allows online selection of the most stable part of thetime-resolved signal. U and Th contents of the ablatedvolumes were estimated by comparison of count rateswith those of the GJ-1 standard, and probably areaccurate to �10%.Hafnium isotopes were analyzed with a Merchantek/

New Wave LUV 213 nm LAM, attached to a NuPlasmamulticollector (MC) ICPMS system; ablations were per-formed in He. Typical ablation pits are 50–60mm indiameter. Procedures for the correction of isotopic inter-ferences of 176Lu and 176Yb on 176Hf and for mass biascorrections have been described by Griffin et al. (2000).Interpretations of Hf isotope evolution utilize the 176Ludecay constant of Blichert-Toft et al. (1997) and the modelfor the depleted-mantle source adopted by Griffin et al.(2000). U–Pb and Hf isotope data on the secondarystandards analyzed as unknowns with each run aresummarized in Table 3.

U–Pb zircon ages

All but four of the 27 analyses of 206Pb/238U and207Pb/235U in zircons yield ages concordant withinanalytical uncertainties in the range from about 85 to60Ma (Fig. 8, Table 4). The 206Pb/238U ages cluster intwo ranges, the larger cluster centered at about 70Ma,

0 100 200 300 400 5000.0

0.1

0.2

0.3

0.4

0.50.5

0.60.6Cr2O3Cr2O3

MnOMnO

TiO2TiO2

0 100100 200200 300300 400400 5005004

8

12

16

20CaOCaO

MgOMgO

FeOFeO

Fe asFeO

CaO

MgO

TiO2

MnO

Cr2O3

Distance (µm) from grain boundary (GB)

Wei

ght p

erce

nt o

xide

(a) ± 2σ

GB

Wei

ght p

erce

nt o

xide

0 1000 2000 3000 4000 50000.0

0.1

0.2

0.3

0.4

Distance (µm)

TiO2

Cr2O3

MnO

(b)

4

8

12

16

20

Fe asFeO

MgO

CaO

paleorange

paleorange

clear

GB?

±2σ

Fig. 4. Compositional zoning documented by EMP analyses. The �2s error bars show the range of 4 SD for a representative analysis, calculatedsolely on the basis of counting statistics. (a) Apparent growth zoning in garnet at a contact with ilmenite in rock N375-GR, along a path markedon the BSE image in Fig. 3a. The grain boundary is marked ‘GB’. (b) Zonation across an area with faint boundaries between a clear central areaand a pale orange halo in rock GR1-202. A probable grain boundary is marked by ‘GB?’.

1907

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

and a smaller one centered at about 85Ma (Fig. 8b).These U–Pb analyses are compatible with a model ofepisodic zircon growth extending over the period fromabout 85 to about 60Ma. The most discordant pointhas a 207Pb/206Pb age of about 670Ma. The discordantpoints document inheritance of a relatively small mass ofzircon 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 grainshave Th/U �0�3 (median 0�49), which might be expec-ted in metamorphic zircons, but others range up to >2.There is no correlation between age and Th/U, but thetwo oldest grains also have the highest Th contents andthe highest Th/U. These older ages may reflect incom-plete Pb loss from originally magmatic grains inheritedfrom the protolith of the garnetite. The high mean con-tents of U and Th are unusual in metamorphic zircons.

Hf isotope ratios and age implications

The analyzed zircons have extraordinarily low176Lu/177Hf and 176Yb/177Hf ratios, comparable with

or lower than those found in kimberlitic zircons (Griffinet al., 2000). The depletion of the HREE is consistent withgrowth in the garnetite matrix, where HREE are stronglypartitioned into garnet, but Hf partitions into zircon.There is no correlation between 176Hf/177Hf and Yb/Hfor Lu/Hf ratios, confirming the efficacy of the isobaric-overlap corrections employed here and described byGriffin et al. (2000).The 176Hf/177Hf determinations extend over the range

0�281829(9)–0�282851(10) (Figs 9 and 10, Table 5). Hfisotope ratios in interiors and rims of selected grains wereobtained both by multiple analyses and by ablatingthrough zircons. There is a crude correlation of176Hf/177Hf with percent HfO2; the zircons with leastradiogenic Hf values have a median of about 2�1 wt %HfO2, and the most radiogenic a median of about1�7 wt %. Also, there is a crude correlation of moreradiogenic Hf with younger 206Pb/238U ages (Fig. 9b).Furthermore, in all cases, grain rims have higher176Hf/177Hf than interiors, and a large part of the totalrange is present within single zircon grains (Fig. 10).Some of the grains were characterized by BSE and cath-odoluminescence (CL) before analysis, and representative

0 300 600 900 1200 15000.0

0.1

0.2

0.3

0.4

0.50.5

0.60.6

TiO2TiO2

300300 600600 900900 12001200 150015004

8

12

16

20CaOCaO

MgOMgO

FeOFeO

Distance (µm)

Wei

ght p

erce

nt o

xide

Fe asFeO

CaO

MgO

MnO

TiO2

Cr2O3

(b)

± 2σ

4

8

12

16

20

CaOMgO

Fe asFeO

0 50 100 150 2000.0

0.1

0.2

0.3

0.4

MnO

TiO2

Distance (µm)

Wei

ght p

erce

nt o

xide

(a)± 2σ

Filledfracture

Fig. 5. Compositional zoning documented by EMP analyses. The �2s error bars show the range of 4 SD for a representative analysis, calculatedsolely 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 theBSE image in Fig. 3c. The traverse increment was 6mm. (b) Complex zoning apparently related both to rimward growth and to fracture fill plusreaction 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 figurewere defined by steps at smaller increments (6mm) than the 36mm steps used to span the remainder of the traverse.

1908

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

Table2:Averagesandranges(ppm

)ofLAM–ICPMSanalysesofgarnetandpyroxene

Sam

ple:

ATG-G

RP1-B

GR1-202

N379b

-GR-4

GR1-201

N379b

-GR-1

Mineral:

Cpx

Gar

Gar

Gar

Gar

Gar*

Gary

Points:

69

82

95

1

Ti

140(100�180)

450(340�550)

420(190�660)

370

270(220�350)

480(380�690)

450

Ni

430(330�570)

11(3�15)

16(11�

27)

8(7,9)

8(4�11)

18(15�

23)

17

Sr

380(340�450)

0.25

(0. 10�

0.34)

0.84

(0. 30�

1.62)

0.10

(0. 07,

0.12)

0.26

(0. 04�

0.44)

0.50

(0. 2�1.0)

0.30

Y0.21

(0. 19�

0.23)

18(10�

22)

16. 8

(15.0�

19. 6)

24(23,

24)

67(52�

93)

12(10�

14)

3

Zr

1.2(1. 0�1.5)

17(9�24)

12. 5

(10.8�

13. 8)

11. 8

(10.5,

13. 1)

2.3(1. 0�2.7)

3.3(2. 2�4.2)

3.4

La

1.5(1. 2�1.9)

b.d.l.

b.d.l.

b.d.l.

b.d.l.

b.d.l.

b.d.l.

Ce

6.6(4. 9�9.0)

0.17

(0. 12�

0.23)

0.41

(0. 20�

0.67)

0.09

(0. 07,

0.10)

0.10

(0. 03�

0.26)

0.07

(0. 06�

0.10)

0.11

Pr

1.6(1. 2�1.8)

0.14

(0. 10�

0.17)

0.25

(0. 16�

0.39)

0.08

(0. 07,

0.09)

0.04

(0. 01�

0.07)

0.05

(0. 04�

0.07)

0.07

Nd

10. 0

(8. 6�11. 9)

2.6(2. 1�3.3)

3.7(3. 1�5.5)

1.6(1. 5,1.8)

0.9(0. 3�1.7)

1.2(0. 9�1.4)

1.6

Sm

1.8(1. 7�2.0)

5.8(2. 1�7.5)

4.9(4. 3�5.8)

2.6(2. 5,2.7)

3.3(1. 7�5.6)

1.9(1. 6�2.3)

2.6

Eu

0.59

(0. 43�

0.69)

1.2(0. 9�1.6)

0.88

(0. 65�

1.13)

1.3(1. 3,1.4)

1.6(0. 9�2.4)

1.5(1. 2�1.7)

1.3

Gd

0.70

(0. 62�

0.81)

4.9(2. 8�6.5)

3.0(2. 6�3.3)

4.5(4. 5,4.5)

6.9(4. 5�10. 1)

2.4(2. 1�2.8)

1.5

Tb

0.05

(0. 03�

0.05)

0.64

(0. 31�

0.88)

0.46

(0. 41�

0.53)

0.65

(0. 64,

0.67)

1.4(1. 0�2.2)

0.40

(0. 35�

0.47)

0.14

Dy

0.10

(0. 07�

0.14)

3.4(1. 6�4.3)

2.7(2. 3�3.2)

4.2(4. 1,4.4)

11. 5

(7. 6�17. 2)

2.5(2. 1�2.9)

0.80

Ho

0.010(0. 01�

0.01)

0.66

(0. 29�

0.84)

0.59

(0. 50�

0.67)

0.87

(0. 81,

0.92)

2.7(2. 1�4.4)

0.44

(0. 36�

0.49)

0.13

Er

0.015(0. 01�

0.02)

1.6(0. 7�2.1)

1.5(1. 3�1.8)

2.6(2. 4,2.7)

8.5(5. 9�13. 4)

1.2(1. 0�1.6)

0.30

Tm

b.d.l.

0.19

(0. 10�

0.25)

0.19

(0. 16�

0.22)

0.40

(0. 36,

0.44)

1.3(0. 8�2.4)

0.16

(0. 13�

0.18)

0.04

Yb

b.d.l.

1.5(0. 7�2.0)

1.5(1. 3�1.8)

3.2(3. 1,3.3)

10. 4

(7. 0�15. 8)

1.4(1. 1�1.6)

0.30

Lu

b.d.l.

0.20

(0. 09�

0.26)

0.22

(0. 17�

0.28)

0.41

(0. 38,

0.44)

1.3(0. 8�2.2)

0.15

(0. 11�

0.19)

0.04

Hf

0.13

(0. 10�

0.14)

0.28

(0. 10�

0.55)

0.15

(0. 12�

0.20)

0.21

(0. 19,

0.23)

0.04

(0. 01�

0.06)

0.03

(0. 01�

0.05)

0.05

b.d.l.,below

determinationlim

it,as

discu

ssed

intext.

*,yIn

terioran

drim,respectively,ofgarnet

(Figs3d

and5b

,an

dTab

le1).

1909

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

images with positions of ablated pits are shown in Fig. 11.Most of the analyzed grains are relatively large and havesmooth shapes (Fig. 11a–c); the interiors of these grainshave relatively old U–Pb ages and relatively low176Hf/177Hf, consistent with the presence of an inheritedProterozoic component in the lased volumes. Zirconsare also present that have botryoidal shapes and granularsurfaces (Fig. 11d and e). Only two grains with thisunusual morphology were large enough to analyze, andthe ablated volumes in each had relatively high176Hf/177Hf (�0�2825). Zircons with similar granularsurfaces were noted in two of the fractions from Navajoeclogites used for U–Pb geochronology by Smith et al.(2004): neither of those multigrain fractions recordedinheritance of Proterozoic lead, in contrast to many ofthe other analyzed fractions from the eclogites. Thebotryoidal morphology may characterize some of thezircon formed in the Cenozoic metamorphic events.The least radiogenic 176Hf/177Hf values yield depleted-

mantle model ages of about 1�9Ga (Fig. 9), and the age ofthe garnetite protolith is likely to have been at least1�8Ga. The depleted-mantle line in Fig. 9 reproducesaverage mid-ocean ridge basalt (MORB) values (Griffinet al., 2000): the range for basalt on the East Pacific Ridge(Chauvel & Blichert-Toft, 2001), plausibly from mantlesources like those for Farallon MORB, is indicated bythe EPR bracket in Fig. 9b. All values of 176Hf/177Hf inthe zircons are less radiogenic than those expected forthe Mesozoic and Cenozoic basalts of the Farallon plate.

Cho

ndrit

e-no

rmal

ized

RE

E a

bund

ance

s

1

100

10

0.1

1

2345

(a)

Sm Gd Dy Er YbCe NdLa Pr Eu Tb Ho Tm Lu57 58 59 60 62 63 64 65 66 67 68 69 70 71

1

100

10

0.1

(c)

1

100

10

0.1

(b)

Garnet

Clinopyroxene

57 58 59 60 62 63 64 65 66 67 68 69 70 71

57 58 59 60 62 63 64 65 66 67 68 69 70 71

Fig. 6. Rare earth element abundances normalized to the C1 chondriteaverages of Sun & McDonough (1989). (a) Average abundances ingarnet 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 rockN379b-GR-1, all within the complexly zoned garnet imaged by BSE inFig. 3d. Five of the analyses cluster closely, as is typical of the analysesfor each of the other rocks plotted in (a). Analyses of the sixth volumeare markedly lower in the HREE; that volume is garnet that has arelatively low average atomic number and so is darkest in the BSEimage (Fig. 3d), and that is relatively lower in Fe (Fig. 5b). (c) Abund-ances in clinopyroxene (dashed lines, six ablations) and garnet(continuous lines, nine ablations) in rock ATG-GRP1-B. The twocurves for garnet with slightly lower Sm abundances were ablated ona later date than the other seven, and the lower abundances of theMREE recorded by the two might be due to either an analyticalproblem or garnet inhomogeneity.

0.4 0.8 1.20

0.04

0.08

0.12

400°C

800°

C50

0°C

600°

C

Fe/Mg garnet

Fe/

Mg

chlo

rite

Pair in garnetpyroxenite

Pair in garnetite

700°

C

Fig. 7. Fe/Mg in garnet and chlorite in Navajo xenoliths in texturesconsistent with equilibrium. Compositions of the five pairs in garnetitesare given in the Electronic Appendix. The three pairs in garnetpyroxenites are from Helmstaedt & Schulze (1988) and Smith (1995).The lines of constant temperature were calculated using the geother-mometer of Dickenson & Hewitt (1986) in the modified form cited byLaird (1988).

1910

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

The more radiogenic 176Hf/177Hf values can be pro-duced by episodic zircon growth in an environmentwith high Lu/Hf. The eight analyses of garnet in sampleGR1-201 have a median Lu/Hf equal to 33, far higherthan the values of 1–2 characteristic of garnet in the othergarnetites. The calculated evolution of 176Hf/177Hf in anenvironment with Lu/Hf of 30 and with an initial valuelike that of the inherited zircon nicely fits the observedcorrelation with 206Pb/238U ages during the period from85 to 70Ma (Fig. 9b).Hf-isotope composition has a weak negative correlation

with U: the grains with the most radiogenic Hf (lowestTDM, where DM indicates depleted mantle) have thelowest U contents. Most low-Th grains also have lowTDM. There is considerable scatter in Th and U contents,however, and no correlation was recognized between Hfisotope composition and Th/U.

PETROLOGICAL FRAMEWORK

AND CONSTRAINTS

Processes of formation of other garnetites

Garnetites, rocks that consist mostly of garnet, havebeen attributed to a variety of processes. For instance,

monomineralic garnet xenoliths in African kimberlitehave been interpreted as magmatic cumulates; the gar-nets contain about 70% pyrope and 10–20% grossularend-members (Exley et al., 1983). Layers of garnetiteconsisting of 60–90 vol. % garnet occur in high-gradeamphibolite- and granulite-facies rocks in the Alps, andthey have been attributed to anatexis and melt–rockinteractions (Rivalenti et al., 1997); garnets in the mostgarnet-rich layers contain about 50% almandine and10% grossular end-members. Garnetite with about20 vol. % corundum occurs in a garnet peridotite lensin the Sulu ultrahigh-pressure terrane and has been inter-preted as a metamorphic rock derived from a protolithof spinel websterite (Zhang et al., 2004). Grossular-richgarnetite boudins in migmatite have been attributed toultrahigh-pressure metamorphism of a subducted calc-silicate protolith (Vrana & Fryda, 2003).The most analogous garnet-rich rocks may be rodin-

gites. Rodingites are formed in metasomatic reactionzones during serpentinization of peridotite in low-temperature near-surface environments, and they occurnear contacts between peridotite and a variety of otherlithologies (Coleman, 1967; Leach & Rodgers, 1978;Schandl et al., 1989; O’Hanley et al., 1992; Dubinskaet al., 2004). Lenses of rodingite typically are less than

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 U�Pb ages (Ma) of standardsy

207Pb/206Pb 2s 207Pb/235U 2s 206Pb/238U 2s 208Pb/232Th 2s

91500 LAM�ICPMS 83 1069.3 3.1 1064.3 3.3 1061.0 4.0 1049 16

TIMS 1065 1

Mud LAM�ICPMS 73 735.5 4.7 731.2 3.2 730.6 3.9 732.4 8.3

Tank TIMS 732 5

Sample No. of analyses Analyses of Hf standardsy

176Hf/ 2s 178Hf/ 2s 180Hf/ 2s 176Yb/ 2s 176Lu/ 2s177Hf 177Hf 177Hf 177Hf 177Hf

JMC475 solution 72 0.282165 13 1.467262 53 1.886774 160

91500 LAM�MC-

ICPMS

634 0.282310 40 1.467043 84 1.886809 222 0.01186 16 0.00032 1

TIMS 0.282290 34 1.467140 24 0.00028 1

Mud LAM�MC- 1477 0.282522 28 1.467127 68 1.886797 166 0.00509 8 0.00011 1

Tank ICPMS

*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 U�Pb, Black & Gulson (1978).

1911

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

30 cm wide and are rarely more than 5m in maximumdimension. The most common garnet in rodingite is fine-grained hydrogrossular; andradite is much less common.Most rodingites contain other minerals, such as diopside,prehnite, zoisite, and tremolite. Temperatures calculatedfrom the mineral assemblages typically are in the range200–500�C, but the higher temperatures may recordmetamorphism after rodingite formation, as describedby Rice (1983). Some instances of the recorded meta-morphism 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 inthe Swiss Alps may be of particular pertinence to theNavajo garnetites, because these Alpine rocks have beeninterpreted to record maximum pressures near 2�5GPa(Evans et al., 1979, 1981). Garnets in the most calciclenses at Cima di Gagnone have grossularite componentsexceeding 75%, but compositions of garnets in rockstransitional between eclogites and metarodingites areless calcic and overlap the range present in the Navajogarnetites.

Temperatures and pressures recordedby Navajo garnetites

Calculations of pressures and temperatures recorded bythe garnetite xenoliths depend upon assumptions thatmineral grains in contact had been in equilibrium. Theranges of garnet composition within some of the smallxenoliths and the sharp compositional boundaries withinsome grains (Fig. 3) are evidence of disequilibrium, as isthe contrast in jadeite component (12 and 25%) betweenthe only two clinopyroxene grains found in garnetiteGR1-203 (Table 1). None the less, an approach toequilibrium is consistent with interelement correlationsbetween samples. For instance, in xenolith N379b-GR-4,garnet, chlorite and ilmenite have the most magnesianand chrome-rich compositions within the suite. The sys-tematic Fe–Mg partitioning between garnet and chloritegrains (Fig. 7) is consistent with equilibrium over a tem-perature range of about 100�C. The partitioning of REEbetween the single garnet–clinopyroxene pair analyzedfor trace elements provides another test. The chondrite-normalized abundances of REE in clinopyroxene form abell-shaped pattern (Fig. 6c). Both the abundance patternand the relative depletion of HREE in the clinopyroxeneare consistent with equilibration with garnet (e.g. Harte &Kirkley, 1997). Because the P–T and compositionaldependences of trace element partitioning betweengarnet and clinopyroxene are poorly known (Blundy &Wood, 2003), quantitative evaluation of equilibriumusing the partitioning is not yet possible for the mineralpair in this xenolith.Compositions used for temperature calculation were

acquired near mutual contacts of chlorite and garnetand of clinopyroxene and garnet. Temperatures werecalculated from Fe/Mg partitioning for chlorite–garnetpairs and Fe2þ/Mg partitioning for clinopyroxene–garnet pairs at a pressure of 2GPa (Table 6); no pressureshave been calculated from the mineral assemblages inthese xenoliths. The thermometer of Krogh (1988) andthe revised thermometer of Krogh Ravna (2000) havebeen applied to garnet–pyroxene pairs; for these appli-cations, ferric iron was calculated in pyroxene from(Na – Al – Cr) in a four-cation formula, and in garnetfrom (2 – Al – Cr) in an eight-cation formula. Temper-atures calculated with the more recent procedure, that ofKrogh Ravna (2000), range between 460�C and 600�Cand average 530�C. Values calculated with the proced-ure of Krogh (1988) are 60–90� higher and average about600�C. The temperatures are not correlated with calcu-lated jadeite or aegirine in pyroxene or with calculatedFe3þ/Fe2þ in garnet (Table 6). Moreover, the correlationbetween calculated Fe3þ/Fe2þ in garnet and pyroxene isevidence that the assumptions used to calculate ferric ironhave validity. Garnet–chlorite temperatures were calcu-lated by two approaches: that of Grambling (1990) and

0.004

0.008

0.012

0.016

0.020

0.024

0.00 0.04 0.08 0.12 0.16 0.20 0.24

160

120

100

80

60

140

40

207Pb / 235U

206 P

b / 2

38U

50 70 90 110 130

Rel

ativ

e pr

obab

ility

206Pb / 238U age (Ma)

(a)

(b)

150

Fig. 8. U–Pb data and ages obtained by LA–ICP-MS analysis ofzircons in garnetite GR1-201. (a) Ratios and concordia plotted usingmethods of Ludwig (2001). Data points are represented by 1s errorellipses. (b) Probability density diagram illustrating the uneven distri-bution of 206Pb/238U ages.

1912

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

Table4:U–Pbanalysesandcalculated

agesofzircon

ingarnetiteGR1-201

Zirco

nTh(ppm)

U(ppm)

Th/U

206Pb/

238U

%RSD

207Pb/

235U

%RSD

207Pb/

206Pb

%RSD

208Pb/

232Th

%RSD

Ages

inMa

206Pb/

1s207Pb/

1s207Pb/

1s208Pb/

1s238U

235U

206Pb

232Th

G201-D

245

872.82

0.01557

1.73

0.11043

4.62

0.05146

4.72

0.00592

1.69

100

2106

5261

74119

2

G201-G

346

0.07

0.01070

2.43

0.04976

15. 39

0.03374

15. 50

0.00200

58. 00

692

497

�125

184

4023

G201-F

1383

0.16

0.01035

2.42

0.06888

7.96

0.04829

8.18

0.00247

15. 38

662

685

114

132

508

G201-N

16126

0.13

0.01302

2.00

0.08390

5.65

0.04676

5.82

0.00355

10. 70

832

824

3785

728

G201-P

33226

0.15

0.01013

1.78

0.06449

4.92

0.04619

5.04

0.00315

7.94

651

633

871

645

G201-M

64150

0.43

0.01093

1.83

0.07204

6.47

0.04782

6.73

0.00347

2.88

701

714

90151

702

G201-H

6298

0.63

0.01262

1.82

0.08476

5.62

0.04874

5.74

0.00408

3.92

811

834

135

9482

3

G201-C

60214

0.28

0.01067

1.78

0.07160

4.69

0.04870

4.83

0.00343

5.25

681

703

133

7569

4

G201-A

135

111

1.22

0.01104

1.90

0.07124

5.99

0.04681

6.15

0.00351

3.13

711

704

4094

712

201-4H

151

267

0.57

0.01376

1.45

0.10448

3.02

0.05508

3.10

0.00527

2.47

881

101

3415

42106

3

201-4G

129

238

0.54

0.01041

1.63

0.06608

3.98

0.04605

4.30

0.00341

3.52

671

653

9169

2

201-4C

1123

436

0.28

0.01056

1.42

0.06688

2.89

0.04595

2.92

0.00389

2.83

67. 7

166

2�5

3378

2

201-4C

273

126

0.58

0.01300

1.62

0.08310

5.48

0.04635

5.72

0.00414

2.66

831

814

16125

842

201-4D

237

982.42

0.02149

1.68

0.18320

3.80

0.06184

3.93

0.00890

1.57

137

2171

6669

53179

3

201-4E

86201

0.43

0.01119

1.61

0.07222

4.36

0.04685

4.46

0.00387

3.62

721

713

4265

783

201-4A

99164

0.60

0.01141

1.58

0.07451

4.27

0.04738

4.35

0.00363

3.03

731

733

6866

732

201-2A

146

239

0.61

0.01147

1.31

0.07414

2.59

0.04689

2.60

0.00396

2.02

73. 5

173

244

3680

2

201-2C

155

175

0.89

0.00981

1.53

0.06452

3.36

0.04771

3.44

0.00253

5.53

62. 9

163

285

5151

3

201-2B

13282

0.05

0.01082

1.48

0.07271

3.16

0.04873

3.22

0.00376

2.39

691

712

135

4776

2

201-2E

13286

0.05

0.01336

1.57

0.09029

3.62

0.04902

3.71

0.00418

2.39

861

883

149

5584

2

201-3B

5298

0.53

0.01087

2.21

0.07195

7.12

0.04800

7.33

0.00282

7.09

702

715

99117

574

201-3E

81166

0.49

0.01109

1.89

0.07327

5.51

0.04790

5.70

0.00305

5.90

711

724

9488

624

201-3G

148

632.35

0.01150

2.87

0.07594

10. 28

0.04787

10. 59

0.00535

3.55

742

747

93175

108

4

201-3H

22117

0.19

0.01213

1.81

0.07967

5.01

0.04764

5.16

0.00325

8.92

781

784

8178

666

201-3K

335

0.09

0.00906

3.31

0.06648

13. 73

0.05323

14. 05

0.00464

24. 14

582

659

339

248

9423

201-3L

28120

0.23

0.01231

1.95

0.08339

5.80

0.04913

5.98

0.00328

9.15

792

815

154

9666

6

201-3N

163

135

1.21

0.01245

1.77

0.07986

6.35

0.04654

6.57

0.00396

2.27

801

785

25145

802

1913

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

that of Dickenson & Hewitt (1986) in the modified formdescribed by Laird (1988). All iron was considered ferrousin both garnet and chlorite for these approaches. Bothchlorite–garnet methods yield very similar temperaturesin the range from 400�C to 500�C, and the average of thefive values is about 470�C. The contrast between thelower garnet–chlorite and the higher garnet–pyroxenetemperatures may be due to at least two causes otherthan disequilibrium—assumptions regarding ferrousiron or faulty calibrations of thermometers for theserocks. The garnet–pyroxene temperatures are sensitiveto assumptions made in calculating ferric and ferrous ironin such magnesian pyroxenes, as emphasized by Proyeret al. (2004). The garnet–chlorite thermometers were cal-ibrated from garnet–biotite thermometry of metamorph-osed pelites in which garnet and chlorite compositionsare unlike those in the garnetites. Hence, the calculatedtemperatures are best regarded only as evidence that the

temperatures of garnetite formation were low, probablyin the range from 400�C to 600�C. Even if inaccurate,the calculated values are meaningful for comparisonsof temperatures calculated by the same methods foreclogite and pyroxenite xenoliths included in the Navajodiatremes.Pressures of garnetite formation are difficult to con-

strain. Comparisons with metamorphosed rodingites inalpine peridotites are complicated by hydration reactionsthat occurred as those rocks were exhumed. For instance,Evans et al. (1979) suggested that metarodingites in anAlpine peridotite had eclogite-facies mineralogies atabout 800�C and 2�5GPa, and that most or all of theamphibole and epidote in those rocks developed during ametamorphic overprint at less than 1GPa. The absenceof amphibole in the garnetites may be a key. Schulze et al.(1987) and Helmstaedt & Schulze (1988) observed thata retrograde assemblage of chlorite–garnet–omphaciteformed as a consequence of hydration of some garnetpyroxenites, whereas pargasite–chlorite formed in others.They concluded that the chlorite–garnet–omphaciteassemblage formed at pressures greater than about2�5GPa, above the stability of amphibole. Neitheramphibole nor epidote was identified in the garnetitexenoliths, despite the evidence for the presence of a fluidphase during garnet growth, and hence the chlorite–garnet–clinopyroxene assemblage in these rocks probablyalso formed near or above about 2�5GPa.

Relevant xenolith assemblages inthe Navajo SUM diatremes

The SUM diatremes in the Navajo volcanic field containunusual xenolith types that provide context for theinterpretation that the garnetites formed in metasomatic

0.2830

0.2828

0.2826

0.2824

0.2820

0.2818

0.2816

0.2822

Outer portionsInteriors

Two or more analyses of individual zircon grains

176 H

f / 1

77H

f

Fig. 10. Values of 176Hf/177Hf in corresponding interior and outerportions of individual zircon grains in garnetite GR1-201. Values forouter portions were obtained both by laser ablation at zircon surfacesnear grain rims and by ablating through grains until garnet wasencountered. The two points representing outer portions of somegrains represent the two types of analysis.

0.2816

0.2820

0.2828

0.2824

0.2832

0 40 80 120

0.2832

0.2812

0.2816

0.2820

0.2824

0.2828

0 0.5 1.0 1.5 2.0

(a)

(b)}

EPR

Zircon (0.000015)

Zircon (0.000015)

DM (0.0384)

CHUR (0.0332)

CHUR (0.0332)

5 My

10 My

15 My

Depleted mantle 176Lu/177Hf = 0.0384

By before present

My before present

176 H

f / 1

77H

f17

6 Hf /

177

Hf

TDM =1.87 Ga

176Lu/177Hf = 4

Fig. 9. 176Hf/177Hf plotted against 206Pb/238U ages from the analysesof zircons in rock GR1-201. Continuous lines and accompanying176Lu/177Hf values model the evolution of depleted mantle (DM) andchondrites (CHUR), as described by Griffin et al. (2000). The bracketlabeled ‘EPR’ spans the range measured for basalts of the East PacificRidge by Chauvel & Blichert-Toft (2001). Dashed lines model theevolution of Hf in zircon with 176Lu/177Hf equal to the measuredvalue of 0�000015 and of 176Lu/177Hf ¼ 4, consistent with theLu/Hf analyses of garnet in the host rock GR1-201.

1914

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

reaction zones. These rock types include those peridotitesand garnet pyroxenites that contain chlorite and otherhydrous minerals and the eclogites. As emphasized byHelmstaedt & Schulze (1988), the diverse lithologies aresimilar to those in some metamorphosed ophiolite com-plexes in high-pressure orogenic belts.Hydrous minerals in spinel peridotite xenoliths

include amphibole, chlorite, titanoclinohumite, and

antigorite: Smith (1979) concluded that all but antigoriteformed by hydration reactions in the mantle, and that atleast some of the antigorite may be of similar mantleorigin. Peridotite xenoliths also contain chlorite attrib-uted to hydration of mantle at greater depths, below thespinel–garnet transition (Mercier, 1976; Smith, 1995).An unusual rock interpreted by Smith (1995) to recordreactions of garnet peridotite with water contains

Table 5: Lu–Hf analyses of zircons, model ages, and corresponding U–Pb ages

Zircon grain 176Hf/177Hf 1s 176Lu/177Hf 176Yb/177Hf TDM (Ga) 206Pb/238U age (Ma)

201/3 B 0.281959 0.000020 0.000020 0.001177 1.71 70

201/3 E interior 0.281852 0.000011 0.000012 0.000668 1.85 71

201/3 E rim 0.282139 0.000019 0.000017 0.000943 1.48 71

201/3 G 0.282186 0.000013 0.000015 0.000739 1.42 74

201/3 H 0.282359 0.000012 0.000015 0.000721 1.19 78

201/3 K-1 0.282581 0.000021 0.000010 0.000541 0.90 58

201/3 K-2 0.282259 0.000015 0.000010 0.000567 1.32 58

201/3 N-1 0.281838 0.000017 0.000015 0.000918 1.87 80

201/3 N-2 0.281915 0.000015 0.000016 0.000966 1.77 80

201/3 N-3 0.282457 0.000058 0.000018 0.001379 1.06 80

201/3 S 0.282738 0.000019 0.000018 0.001055 0.69

201/4 A 0.281973 0.000015 0.000016 0.000910 1.70 73

201/4 A rim 0.282413 0.000017 0.000011 0.000563 1.12 73

201/4 C-1 interior 0.281903 0.000011 0.000015 0.000960 1.79 66

201/4 C-1 rim 0.282400 0.000020 0.000015 0.000960 1.13 66

201/4 C-2 interior 0.282196 0.000015 0.000016 0.000955 1.40 81

201/4 C-2 rim 0.282454 0.000025 0.000016 0.000880 1.06

201/4 C-2 rim 0.282434 0.000012 0.000012 0.000701 1.09

201/4 D 0.282710 0.000023 0.000015 0.000832 0.72 137

201/4 G interior 0.282070 0.000016 0.000013 0.000679 1.57 67

201/4 G rim 0.282437 0.000015 0.000013 0.000704 1.09 67

201/4 H 0.281883 0.000010 0.000014 0.000770 1.81 88

201 A interior 0.281886 0.000018 0.000019 0.001047 1.81 71

201 A rim 0.282651 0.000024 0.000028 0.001400 0.80 71

201 E 0.281829 0.000009 1.88

201 D 0.281862 0.000006 0.000012 0.000694 1.84 100

201 G interior 0.282456 0.000017 0.000005 0.000210 1.06 69

201 G rim 0.282567 0.000021 0.000010 0.000430 0.91 69

201 F 0.282851 0.000010 0.000011 0.000536 0.54 66

201 N 0.281985 0.000011 0.000017 0.000930 1.68 83

201 P 0.282140 0.000021 0.000014 0.000795 1.48 65

201M 0.282305 0.000015 1.26 70

201 H interior 0.281949 0.000017 0.000014 0.000836 1.73 81

201 H rim 0.282382 0.000025 0.000017 0.000832 1.16

201 H rim 0.282844 0.000011 0.000013 0.000651 0.55

201 C 0.281911 0.000009 0.000015 0.000851 1.78 68

201/2 A 0.282622 0.000021 0.000023 0.001058 0.84 74

201/2 C 0.282132 0.000013 0.000017 0.000940 1.49 63

201/2 B 0.281868 0.000011 0.000015 0.000724 1.83 69

201/2 E 0.281946 0.000011 0.000019 0.001174 1.73 86

interior, interior of zircon grain; rim, outer portion of zircon grain.

1915

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

centimeter-scale volumes with contrasting proportions ofchlorite, clinopyroxene, orthopyroxene, ilmenite, andtitanian chondrodite. Minerals in these volumes appearto have formed in metasomatic reaction zones, consistent

with element transport analogous to that during rodingiteformation.Chlorite appears together with a second generation of

garnet and clinopyroxene in some garnet pyroxenite

60 µm

60 µm 50 µm60 µm

0.2819

0.2819

0.2824,0.2828

68 Ma81 Ma

0.2819 100 Ma

0.2825,0.2826

69 Ma

100 µm

(a)

(b)(c)

(d) (e)

Fig. 11. Images of zircons in garnetite GR1-201 polished thin sections, together with positions of laser pits and corresponding 176Hf/177Hf ratiosand 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

Xenolith T (�C) T (�C) % jadeite % aegirine Fe3þ/Fe2þ Fe3þ/Fe2þ T (�C) T (�C)

Cpx�Gar1 Cpx�Gar2 in Cpx3 in Cpx4 in Cpx in Gar Chl�Gar5 Chl�Gar6

GR-41 549 492 14 6 1.1 0.05 —— ——

N379b-GR-2 528 461 5 2 0.31 0.02 —— ——

N379b-GR-3 690 602 6 1 0.14 0.01 —— ——

GR1-2037 618 548 12 5 0.8 0.05 —— ——

GR1-2037 606 531 25 6 1.4 0.05 402 411

ATG-GRP1-B 631 541 4 1 0.13 0.03 492 497

N379b-GR-4 —— —— —— —— —— —— 497 501

GR-P1-ATG —— —— —— —— —— —— 431 441

N375-GR —— —— —— —— —— —— 492 497

1Procedure of Krogh (1988) with ferric iron calculated as discussed in text.2Procedure of Krogh Ravna (2000) with ferric iron calculated as discussed in text.3Calculated equal to total Al in a four-cation formula.4Calculated equal to (Na�Al�Cr) in a four-cation formula.5Procedure of Dickenson & Hewitt (1986) in form published by Laird (1988).6Procedure of Grambling (1990).7The two clinopyroxene grains present are compositionally distinct.

1916

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

xenoliths (Helmstaedt & Schulze, 1979; Benoit &Mercier, 1986; Schulze et al., 1987; Smith, 1995).Chlorite–garnet thermometry records temperatures inthe range from 400�C to 500�C for the pairs in apparenttextural equilibrium (Smith, 1995), just as recorded bygarnet–chlorite pairs in the garnetites (Fig. 7). Garnetsformed with chlorite in the pyroxenites have com-positions near Gr28Prp36Alm36, similar to those in somegarnetites (Fig. 2), as also noted by Helmstaedt & Schulze(1988). However, clinopyroxene in the retrogradeassemblages of the garnet pyroxenites has 37–47% Jdcomponent, more sodic than the range from 3% to25% found for pyroxene in the garnetites.Eclogite xenoliths record garnet–pyroxene temperat-

ures between about 500�C and 700�C (e.g. Helmstaedt& Schulze, 1988; Smith et al., 2004), similar to the valuescalculated for the garnetites (Table 6). Coesite has beenidentified in one eclogite, establishing a minimum pres-sure of about 2�5GPa (Usui et al., 2003). Three eclogitesrecord pressures in the range 2�6–3�4GPa by garnet–phengite–omphacite barometry, and the calculated pres-sures and temperatures are consistent with the presenceof lawsonite in these rocks (Smith et al., 2004). Textures ofthe eclogites indicate the presence of a water-rich fluidduring recrystallization (Smith & Zientek, 1979), andthe formation of sodic eclogite and associated jadeiteand omphacite pyroxenite has been attributed to hydrousmetasomatism (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 ofwhich appears primary, occurs in five of the xenoliths,and these rocks are otherwise typical of the suite. Tem-peratures calculated from garnet–chlorite and garnet–clinopyroxene pairs by all methods are in the range400–700�C, and those calculated by the more recentmethods are in the range 400–600�C (Table 6). Manygarnets contain compositional gradients that are sharp ona scale of several micrometers, as documented by theback-scattered electron images in Fig. 3 and by the tra-verses plotted in Fig. 5. Such compositional gradientswould have at least partly annealed, if cooling had beenslow enough to reset the garnet–clinopyroxene andgarnet–chlorite temperatures. Therefore, the mineralogy,geothermometry, and textures are metamorphic. Second,although the temperature constraints could be consistentwith low-temperature metamorphism of an igneous pro-tolith, the xenolith compositions are unlike those of therare garnetites attributed to igneous processes. The igne-ous garnetites of mantle origin described by Exley et al.(1983) and Zhang et al. (2004) are more magnesian thanthe xenoliths. Those interpreted as deep-crustal restites

by Rivalenti et al. (1997) contain less calcic garnet. Thusthere is no evidence that the unusual bulk compositions ofthe Navajo garnetites formed by an igneous process.The bulk compositions, mineralogies, and textures of

the garnetites instead are attributed to formation in meta-somatic reaction zones at contacts of mafic rock andperidotite within the continental mantle lithosphere.The reactions are inferred to be a consequence of fluidflow from hydrating peridotite into the mafic rock,a process analogous to that which forms rodingite.Rocks that are monomineralic, or nearly so, are com-monly formed in metasomatic zones (Thompson, 1959).Coleman (1967) and many others have described rodin-gites that formed by metasomatism at contacts of serpent-inized peridotite. The fracture-related garnet growthestablishes that a hydrous fluid was present during gar-netite formation, as required for the process. Althoughthe garnets are unlike those of rodingites, the differencesmay be due to relatively higher temperatures and pres-sures of garnetite formation. REE patterns of the garnetsare also consistent with formation in metasomatic reac-tion zones. The patterns are not like those for manymantle garnets, in which ‘convex-up’ chondrite-normal-ized plots decrease smoothly from Lu to La. Rather, infour of the five rocks, abundances of the middle REE(MREE) are relatively high, more similar to the‘sinusoidal’ or ‘humped’ patterns interpreted as con-sequences of metasomatism by Hoal et al. (1994), Roden& Shimizu (2000), Burgess & Harte (2004), and Zhanget al. (2004).Garnetite formation during serpentinization is consist-

ent with the stability of antigorite in the mantle. Alumin-ous antigorite is stable to temperatures above 660�C at2GPa (Bromiley & Pawley, 2003), hotter than almost alltemperatures calculated for the garnetites (Table 6).Bromiley & Pawley (2003) also demonstrated that anti-gorite with 3�1 wt % Al2O3 is stabilized to higher tem-peratures than is pure antigorite, and antigorite in Navajoperidotite xenoliths contains as much as 3�9 wt % Al2O3

and 1�2 wt % Cr2O3 (Smith, 1979).Compositions of the minor minerals are consistent

with the hypothesis that peridotite was involved in themetasomatic process. Clinopyroxene has calculatedFe2þ/(Fe2þ þ Mg) in the range 0�05–0�08. Chlorite andphlogopite have Fe/(Fe þ Mg) of 0�05–0�08 and 0�10,respectively. These ratios are similar to those in someperidotites. The minor minerals also have appropriateNi contents. Analyzed chlorite has NiO in the range0�24–0�44 wt %, with an average of 0�34 wt %. Forcomparison, chlorite in equilibrium with olivine in alpineperidotites has NiO in the range 0�19–0�23 wt % in therocks studied by Trommsdorff & Evans (1969, 1974) andSmith (1979).The ranges of mineral compositions and the varieties of

compositional zoning within garnet may record sample

1917

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

positions within individual reaction zones and the evolu-tion of the zones with time. MgO and Cr2O3 are zoned tohigher values at rims of some garnets (Figs 4a and 5b),consistent with changes expected as a mafic lithology isinfiltrated 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 inone xenolith (Fig. 5a) and relatively high in another(Fig. 5b). The ranges in trace element composition maybe due to the presence of gradients of chemical potentialat a variety of scales, as discussed for metarodingiteformation by Frost (1975): Zr concentrations in garnet(Table 2) range from a low of 2 ppm in the rock withrelatively abundant zircon (GR1-201) to a high of 17 ppmin a rock in which zircon was not found (ATG-GRP1-B).Rodingite formation commonly is interpreted as a pro-

cess that occurs in the upper crust (e.g. Evans, 1977), butthe ages deduced from U–Pb and Hf isotopic data insteadare compatible with geochronology of the Plateau man-tle. U–Pb ages establish that much of zircon in garnetiteGR1-201 formed from 85 to 60Ma, but no significantPhanerozoic thermal event has been recognized in geo-chronological studies of crustal xenoliths in the Navajodiatremes (Condie et al., 1999; Selverstone et al., 1999;Crowley et al., 2004). Most of the age range recordedby concordant garnetite zircon is within the 81–33Maperiod of growth of concordant zircon in the associatedeclogite xenoliths (Usui et al., 2003; Smith et al., 2004),and eclogite recrystallization has been attributed to man-tle processes, either in the subducted Farallon slab (Usuiet al., 2003) or in the overlying mantle wedge (Wendlandtet al., 1996; Smith et al., 2004). Garnetite genesis withinthe Farallon slab is precluded by the U–Pb and Hfisotope data that establish inheritance of Proterozoiczircon, but the data are consistent with garnetite forma-tion in the Proterozoic mantle wedge during Farallonsubduction.

EVOLUTION OF THE MANTLE

BELOW THE DIATREME

Age and stability of lithosphere

The inherited zircon in garnetite GR1-201 adds animportant constraint for the evolution of the lithosphereof the Colorado Plateau. The eclogites with the mostbasalt-like compositions have Nd model ages of 2�7Ga(Roden et al., 1990) and 1�5–1�8Ga (Wendlandt et al.,1993), but interpretation of these ages is hindered by theREE metasomatism that affected the suite. Re-depletionOs model ages of peridotite xenoliths from a minette plugin the Navajo field range from 1�1 to 1�8Ga; the rangemay be due to addition or loss of Re before eruption (Leeet al., 2001). The Hf depleted mantle model age of at least1�8Ga for garnetite zircon is a more robust indicator of

when the Plateau mantle was stabilized. That age ismatched by the oldest Hf depleted mantle model agesof zircon cores in xenoliths from the lower crust, also1�8Ga (Crowley et al., 2004), and it is consistent withNd model ages of crustal xenoliths that establish themain crust-forming event at about 1�85Ga (Wendlandtet al., 1993). The correspondence of ages for mantle andcrustal formation is evidence that the crust and upper-most mantle have been coupled together since initialcrustal formation.

Thermal histories and a tectonic model

Calculated temperature histories provide insights into thedepth of garnetite formation. A possible pre-Laramidegeotherm (Fig. 12) was calculated for a lithosphere of200 km thickness that has a basal temperature of 1300�Cand a surface heat flow of 56 mW/m2; the heat flow is theaverage of the two values closest to the diatreme in thecompilation of Minier & Reiter (1991). Geotherms alsoare plotted for a model of the lithosphere during flatsubduction of the Farallon slab, 700 km from the trenchwith a relative plate convergence of 10 cm/year. Thegeotherms for the period during Farallon subductionwere calculated using a standard finite difference con-ductive thermal model similar to but simpler than thatused by Spencer (1996). He calculated geotherms formodels in which the Farallon slab sheared off andreplaced the lower part of the Plateau lithosphere: themost significant difference in the model used here is thatthe geotherm in the Farallon plate at the trench was notadjusted for the age of the subducting slab. The interplate

0

50

100

150

km

0 200 400 600 800 1000°C

Crust

Mantle lithosphere

Farallon slab

010

4020

2

3

1Antigorite out

a

Fig. 12. Schematic illustration of simplified models of the lithospherebelow the Garnet Ridge diatreme. Dashed lines show calculatedgeotherms. That labeled ‘0’ is a steady-state geotherm calculated torepresent the lithosphere before the beginning of subduction. Thethree other geotherms are for 10, 20, and 40Myr after the base ofthe lithosphere has been truncated at 150 km depth by the Farallonslab during flat subduction. The bold continuous line represents thehigh-temperature stability limit of aluminous antigorite determined byBromiley & Pawley (2003). The shaded area around point 1 is appro-priate for the formation of garnetite sample GR1-201, as constrainedby the concordant U–Pb zircon ages from 85 to 60Ma. Points 2 and 3represent eclogite xenoliths with dated zircons, as discussed in the text.

1918

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

contact was assumed to be at 150 km: this thickness forthe remnant lithosphere is consistent with the minimumvalue of 120 km based on analyses of Os in xenolithserupted at about 25Ma (Lee et al., 2001) and with thevalue of 120–150 km based on interpretations of datafrom the LA RISTRA seismic line (West et al., 2004). TheLA RISTRA array passes about 5 km from the diatreme.The crustal thickness of about 47 km is constrained byreceiver function analysis of data from that same seismicarray (Wilson et al., 2003). The model used for geothermcalculation is simplified and the necessary assumptionsare not well constrained, but one conclusion is relativelyrobust: before and near the beginning of low-angle sub-duction, and if temperatures were determined only byconductive heat flow, then antigorite was stable only inthe upper few tens of kilometers of the continental mantlelithosphere.The oldest nearly concordant U–Pb ages of zircon

garnetite xenolith GR1-201 cluster near 85Ma (Table 4,Fig. 8). In contrast, flat subduction has been inferred tohave occurred during the Laramide orogeny, from about80 to 40Ma (Coney & Reynolds, 1977; Spencer, 1996).Geological relationships on the Colorado Plateau also areconsistent with the beginning of the Laramide orogeny atabout 80Ma. Intrusions in the Carrizo Mountains, about50 km east of the Garnet Ridge diatreme, were emplacedfrom 74 to 71Ma and have been ascribed to Laramideprocesses (Semken & McIntosh, 1997). In northern NewMexico east of the Carrizo Mountains, stratigraphic andstructural effects of the Laramide orogeny began in theperiod from 80 to 75Ma (Cather, 2004). If the modellithosphere and geotherms in Fig. 12 reproduce condi-tions before and during the Laramide orogeny, then attimes before and within about 15Myr after the beginningof flat subduction, mantle serpentinization could occuronly at depths shallower than about 85 km. If so, and ifthe garnetites are rodingite analogues, then they musthave formed in the uppermost mantle.The actual history of the continental lithosphere below

the diatremes during Farallon subduction may have beenmuch more complex than that assumed for constructionof Fig. 12. Complexities in the mantle today are illustratedin a model derived from compressional and shear seismicphases recorded by the LA RISTRA seismic array (Gaoet al., 2004). Their tomographic reconstructions havesteeply dipping boundaries in seismic velocity that extendbeneath the Navajo volcanic field from 50 km to 200 kmdepth, and they suggest that these boundaries are deter-mined by temperature differences related to Proterozoicstructures. The presence of a mantle suture below thediatreme is consistent with the suggestion by Selverstoneet al. (1999) that the eclogite xenoliths are from a Prot-erozoic subduction zone. Fluid infiltration and serpent-inization above subducting slabs may be controlled bystructural features (Hyndman & Peacock, 2003). The rise

of buoyant serpentinite masses, such as those discussedby Guillot et al. (2001) and Ueda et al. (2004), could alsobe controlled by existing structures. The conductive geo-therms plotted in Fig. 12 provide only a starting point toconsider depths of origin of the garnetite xenoliths, as thetemperatures in and near the probable suture zone mayhave been influenced by the flow of serpentinite and ofhydrous fluids.

Relevance to the genesis of associatedeclogites

Formation of the garnetites in metasomatic reactionzones demands the presence of one or more other rocktypes to react with peridotite, and the eclogite xenolithsrepresent possible reactants. Both garnetite GR1-201 andsome of the eclogites inherited zircon from Proterozoicprotoliths and also contain zircon that grew during theinterval from 85 to 33Ma (Usui et al., 2003; Smith et al.,2004). Both rock types retain evidence of protoliths withlow-pressure histories. The eclogites probably representsubducted oceanic crust, because of their oxygen isotopecompositions (Smith et al., 2004), and because a negativeEu anomaly in one eclogite has been attributed to pla-gioclase fractionation (Roden et al., 1990). The chondrite-normalized REE pattern (Fig. 6) for one garnetite has apositive Eu anomaly that also may be inherited fromsubducted oceanic crust. Although the eclogites containmore jadeite-rich clinopyroxene than do the garnetites,the contrast is like that between pyroxenes in geneticallyrelated eclogites and metarodingites described by Evanset al. (1979, 1981).The possibility that garnetite and eclogite xenoliths

were from the same mantle source region can be testedby comparisons of pressures and temperatures recordedby two Navajo eclogites that contain dated zircons. Pres-sures and temperatures calculated by Smith et al. (2004)for these two rocks yield depths greater than those out-lined for garnetite formation in Fig. 12. One of the xeno-liths contains a fraction of concordant zircon that grewat about 39Ma (Smith et al., 2004), near the end of theLaramide orogeny, and about 10Myr before diatremeeruption. That eclogite plots at a temperature slightlyhigher than the boundary determined for antigorite sta-bility by Bromiley & Pawley (2003) (Fig. 12, point 3). Theother eclogite contains zircon that yielded concordantages from 81 to 47Ma (Usui et al., 2003), and it plots atconditions too cool for any of the calculated geotherms,even after 40Myr of low-angle subduction (Fig. 12,point 2). The thermobarometry may be inaccurate, andthe model may not reproduce the temperature evolutionof the mantle. If so, the calculations and the model donot preclude a common source region for the xenoliths,despite these discrepancies. Regardless, the similar zirconchronologies and cool garnet–pyroxene temperatures are

1919

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

evidence that garnetite and eclogite record a commonprocess related to rock–fluid interactions.

Hydration of the mantle lithosphere

Formation of the garnetites as rodingite analogues inProterozoic mantle demands the presence of water, andat least two hydration episodes are recorded by the xeno-liths and xenocrysts. Pyrope grains are scattered through-out the SUM host rock at Garnet Ridge, and some ofthese garnets contain inclusions of hydrous minerals suchas chlorite, titanoclinohumite, and carmichaelite as wellas of olivine and pyroxene (McGetchin & Silver, 1970;Hunter & Smith, 1981; Wang et al., 1999). The pyropegrains with included chlorite have complex REE patternssimilar to some of those plotted in Fig. 6 (Roden &Shimizu, 2000). The included hydrous minerals appearto have been trapped in pyrope during prograde garnetgrowth. In contrast, chlorite, amphibole, antigorite, andother hydrous minerals also formed during retrogradehydration of spinel and garnet peridotite and pyroxenite(Mercier, 1976; Helmstaedt & Schulze, 1979; Smith,1979, 1995). The retrograde hydration is more plausiblyrelated to water introduced into the mantle wedgeduring low-angle Farallon subduction accompanyingthe Laramide orogeny. In is unclear which hydrationevent, if either, is related to garnetite formation and thegrowth of concordant zircon in garnetite and eclogite atabout 80Ma.The volume of mantle hydrated during the Laramide

orogeny cannot be constrained from these xenoliths. Thehydrated sources of the SUM diatremes must have beendeeper than those of the xenoliths, and so below about110 km, if the calculated depths and temperatures plottedin Fig. 12 are accurate. Smith et al. (2004) suggested thatthe hydration might have been restricted to long-livedtectonic boundaries within the mantle, like those dis-cussed by Selverstone et al. (1999) and Gao et al. (2004).In contrast, Humphreys et al. (2003) suggested thatdehydration of the Farallon slab resulted in such extens-ive hydration of the overlying mantle that it causedregional uplift of the western USA. Regardless of theextent of hydration, the garnetites provide evidence ofat least local hydrous metasomatism in the continentalmantle near the beginning of the Laramide orogeny.If the hydration was due to water released from theFarallon slab or mobilized by associated magmatism,then the oldest concordant U–Pb ages of about 85Mamay be useful in constraining the history of that slab.

DISCUSSION

Timing and duration of metasomatism

The concordant zircon U–Pb ages in sample GR1-201and in the eclogite samples studied by Usui et al. (2003)

and Smith et al. (2004) scatter over a period of at least30Myr. However, the presence of inherited componentsof clearly older ages makes the U–Pb data ambiguous interms of the timing of the metasomatic events that pro-duced the eclogites and garnetites, because any givengrain (or a given ablated volume) might contain a mixtureof inherited and metamorphic zircon, giving mixed ages.U–Pb data alone also cannot distinguish between thegrowth of new zircon and complete loss of Pb fromolder zircons. Zheng et al. (2004a, 2004b) have demon-strated howHf 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 avail-able from U–Pb data, and they provide a rough estimateof the duration of the metasomatic event(s). The widespread in 176Hf/177Hf is consistent with the rapid evolu-tion of radiogenic Hf in the garnetite matrix (Fig. 9).Scherer et al. (1997) suggested that the presence of zirconwould buffer the Lu/Hf of a garnet granulite to lowvalues. This clearly was not the case in garnetite GR1-201, as Hf has been sequestered into zircon and Lu intogarnet, leaving the garnet matrix with an extremely highLu/Hf and the zircon with an extremely low Lu/Hf. Themigration of radiogenic Hf from the garnet into existingzircons (producing the observed Hf-isotope zoning), andits incorporation into several generations of newly grownzircon, would be enhanced by the periodic movement offluids through the matrix that is suggested by the major-element zoning within garnet (Figs 3 and 5).The 15Myr required for the evolution of the observed

spread in Hf isotope ratios (Fig. 9) may be a minimumestimate for the time over which episodic metasomatismand zircon growth took place within the garnetite. Someof the scatter about the growth curve (Fig. 9) may reflectthe mixing of domains with different ages and differentHf-isotope compositions, at the scale of the laser-ablationanalyses. Assuming that the most radiogenic Hf-isotopecompositions have not been measured because of mixingof domains during ablation, the metasomatic episodesmay have spread over a longer time, such as the 20–25Myr reflected by the main spread of zircon ages(Fig. 8). However, it seems clear that the metasomatismthat produced the garnetites occurred over a short period(15–25Myr) near the end of the Cretaceous and thebeginning of the Laramide orogeny, and that garnetiteformation overlapped the period of formation of somezircons in the associated eclogites.

Implications for processes in themantle wedge

The garnetite provides unusual evidence for rock–waterinteractions in continental mantle. Except for examplesfrom the Colorado Plateau, most xenolith evidence forsuch hydration is restricted to the presence of amphibole

1920

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

and phlogopite, as summarized by Luth (2004). Eventhough serpentinization of continental mantle may be acommon and important process in the forearc wedge(Hyndman & Peacock, 2003), most evidence for it isbased on seismic data. The hypothesis that low-anglesubduction has caused extensive and widespread hydra-tion of continental lithosphere (Humphreys et al., 2003) isdifficult to test. If the garnetites are rodingite analogues,then they document serpentinization above the Farallonslab and at about 700 km from the trench, consistent withthat hypothesis. The zircon U–Pb ages establish thatgarnetite formation began 5–10Myr before Laramidetectonism, however, and the timing may indicate thathydration and low-angle subduction can precede relatedtectonism and magmatism by that time interval. Calcu-lated pressures and temperatures together with zirconU–Pb ages do not fit simple thermal histories of conduct-ive heat transfer in the mantle above the Farallon slab(Fig. 12). Coupled geochronological and petrologicalanalyses of additional xenoliths from the Navajo dia-tremes may document how fluid flow and serpentinediapirism can affect the mantle wedge.The garnetite xenoliths also provide evidence of mobil-

ity of elements that are important in deciphering mech-anisms of arc volcanism. Some of these elements, such asTi, Zr, and Hf, have been termed ‘conservative’ andimmobile in aqueous fluids (Pearce & Peate, 1995), par-ticularly in the presence of rutile (Brenan et al., 1994).Evans et al. (1981) observed that Ti, Zr, and Hf appearedto have been immobile in formation of metarodingites,but such immobility may not be the rule at either crustor mantle pressures. Zr and Ti can be mobile duringrodingite formation at low pressures (Dubinska et al.,2004) and at relatively high pressures during formationof eclogite veins containing zircon and rutile (Philippot &Selverstone, 1991; Rubatto & Hermann, 2003).Woodhead et al. (2001) found evidence for nonconservat-ive behavior of Hf in slab–wedge interactions, perhapsbecause of fluid transport of Hf, a suggestion they viewedas radical. The Hf isotope abundances and the U–Pbages testify to the mobility of Zr and Hf in formation ofgarnetite GR1-201.

ACKNOWLEDGEMENTS

Some of the garnetite samples were collected byW. C. Hunter and by A. T. Gavasci. LAM–ICPMSanalyses at The University of Texas at Austin wereacquired with the assistance of J. Lansdown. NormPearson and Suzie Elhlou are thanked for their cheerfuland patient assistance with the U–Pb and Lu–Hf ana-lyses. For many years, D. J. Schulze has shared informa-tion about Navajo garnetite and eclogite xenoliths, andinsights into their petrogenesis, much to the benefit ofthis contribution. M. F. Roden made valuable comments

during manuscript preparation and on a preliminarydraft. This paper was much improved as a result ofconstructive reviews by J. Selverstone, J. G. Liou andC. Shaw, and by editor B. R. Frost. This is publication383 from the ARC National Key Centre for Geo-chemical Evolution and Metallogeny of Continents(www.es.mq.edu.au/GEMOC). Support for research wasprovided by the Geology Foundation of The Universityof Texas at Austin.

SUPPLEMENTARY DATA

Supplementary data for this paper are available at Journalof Petrology online.

REFERENCES

Belousova, E. A., Griffin, W. L., Shee, S. R., Jackson, S. E. &

O’Reilly, S. Y. (2001). Two age populations of zircons from the

Timber Creek kimberlites, Northern Territory, Australia as

determined by laser ablation ICPMS analysis. Australian Journal of

Earth Sciences 48, 756–766.

Benoit, V. & Mercier, J. C. (1986). Les enclaves ultramafiques du

volcanisme alcalin tertiaire du centre du plateau du Colorado:

implications tectoniques.Bulletin de la Societe Geologique de France 6, 1015–

1023.

Black, L. P. & Gulson, B. L. (1978). The age of the Mud Tank

carbonatite, Strangways Range, Northern Territory. BMR Journal of

Australian Geology and Geophysics 3, 227–232.

Blichert-Toft, J., Chauvel, C. & Albarede, F. (1997). The Lu–Hf

geochemistry of chondrites and the evolution of the mantle–crust

system. Earth and Planetary Science Letters 148, 243–258.

Blundy, J. & Wood, B. (2003). Partitioning of trace elements between

crystals and melts. Earth and Planetary Science Letters 210, 383–397.

Brenan, J. M., Shaw, H. F., Phinney, D. L. & Ryerson, F. J. (1994).

Rutile–aqueous fluid partitioning of Nb, Ta, Hf, Zr, U and Th:

implications for high field strength element depletions in island-arc

basalts. Earth and Planetary Science Letters 128, 327–339.

Bromiley, G. D. & Pawley, A. R. (2003). The stability of antigorite in

the systems MgO–SiO2–H2O (MSH) and MgO–Al2O3–SiO2–H2O

(MASH): the effects of Al3þ substitution on high-pressure stability.

American Mineralogist 88, 99–108.

Burgess, S. R. & Harte, B. (2004). Tracing lithosphere evolution

through the analysis of heterogeneous G9–G10 garnets in peridotite

xenoliths, II: REE chemistry. Journal of Petrology 45, 609–634.

Cather, S. M. (2004). Laramide orogeny in central and northern New

Mexico and southern Colorado. In: Mack, G. H. & Giles, K. A. (eds)

The Geology of New Mexico. New Mexico Geological Society Special Publication

11, 203–248.

Chauvel, C. & Blichert-Toft, J. (2001). A hafnium isotope and trace

element perspective on melting of the depleted mantle. Earth and

Planetary Science Letters 190, 137–151.

Coleman, R. G. (1967). Low-temperature reaction zones and alpine

ultramafic rocks of California, Oregon and Washington. US

Geological Survey Bulletin 1247, 49 pp.

Condie, K. C., Latysh, N., Van Schmus, W. R., Kozuch, M. &

Selverstone, J. (1999). Geochemistry, Nd and Sr isotopes, and U/Pb

zircon ages of granitoid and metasedimentary xenoliths from the

Navajo volcanic field, Four Corners area, southwestern United

States. Chemical Geology 156, 95–133.

1921

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

Coney, P. J. & Reynolds, S. J. (1977). Cordilleran Benioff zones. Nature

270, 403–406.

Crowley, J. L., Schmitz, M. D., Bowring, S. A., Williams, M. L. &

Karlstrom, K. E. (2004). U–Pb and Hf isotopic analysis of zircon in

lower crustal xenoliths from the Navajo volcanic field: 1�4Ga mafic

magmatism and metamorphism beneath the Colorado Plateau.

Geological Society of America, Abstracts with Programs 36, 47.

Dickenson, M. P. & Hewitt, D. A. (1986). A garnet–chlorite

geothermometer. Geological Society of America, Abstracts with Programs

18, 584.

Dubinska, E., Bylina, P., Kozlowski, A., Dorr, W., Nejbert, K.,

Schastok, J. & Kulicki, C. (2004). U–Pb dating of serpentinization:

hydrothermal zircon from a metasomatic rodingite shell (Sudetic

ophiolite, SW Poland). Chemical Geology 203, 183–203.

Evans, B. W. (1977). Metamorphism of alpine peridotite and

serpentinite. Annual Review of Earth and Planetary Sciences 5, 397–447.

Evans, B. W., Trommsdorff, V. & Richter, W. (1979). Petrology

of an eclogite–metarodingite suite at Cima di Gagnone, Ticino,

Switzerland. American Mineralogist 64, 15–31.

Evans, B. W., Trommsdorff, V. & Goles, G. G. (1981). Geochemistry

of high-grade eclogites and metarodingites from the central Alps.

Contributions to Mineralogy and Petrology 76, 301–311.

Exley, R. A., Smith, J. V. & Dawson, J. B. (1983). Alkremite, garnetite

and eclogite xenoliths from Bellsbank and Jagersfontein, South

Africa. American Mineralogist 68, 512–516.

Frost, B. R. (1975). Contact metamorphism of serpentinite, chloritic

blackwall and rodingite at Paddy-Go-Easy Pass, central Cascades,

Washington. Journal of Petrology 16, 272–313.

Gao, W., Grand, S. P., Baldridge, W. S., Wilson, D., West, M., Ni, J. F.

& Aster, R. (2004). Upper mantle convection beneath the central

Rio Grande Rift imaged by P and S wave tomography. Journal of

Geophysical Research 109, B03305, doi:10.1029/2003JB002743.

Grambling, J. A. (1990). Internally-consistent geothermometry and

H2O barometry in metamorphic rocks: the example garnet–

chlorite–quartz. Contributions to Mineralogy and Petrology 105, 617–628.

Griffin, W. L., Pearson, N. J., Belousova, E., Jackson, S. E.,

O’Reilly, S. Y., van Achterberg, E. & Shee, S. R. (2000). The Hf

isotope composition of cratonic mantle: LAM–MC-ICPMS analysis

of zircon megacrysts in kimberlites. Geochimica et Cosmochimica Acta

64, 133–147.

Guillot, S., Hattori, K. H., de Sigoyer, J., Nagler, T. & Auzende, A.-L.

(2001). Evidence of hydration of the mantle wedge and its role in

the exhumation of eclogites. Earth and Planetary Science Letters 193,

115–127.

Harte, B. & Kirkley, M. B. (1997). Partitioning of trace elements

between clinopyroxene and garnet: data from mantle eclogites.

Chemical Geology 136, 1–24.

Helmstaedt, H. H. & Doig, R. (1975). Eclogite nodules from

kimberlite pipes of the Colorado Plateau—samples of subducted

Franciscan-type oceanic lithosphere. Physics and Chemistry of the Earth

9, 95–111.

Helmstaedt, H. H. & Schulze, D. J. (1979). Garnet clinopyroxenite–

chlorite eclogite transition in a xenolith from Moses Rock: further

evidence for metamorphosed ophiolites under the Colorado Plateau.

In: Boyd, F. R. & Meyer, H. O. A. (eds) The Mantle Sample: Inclusions

in Kimberlites and Other Volcanics. Washington, DC: American

Geophysical Union, pp. 357–365.

Helmstaedt, H. H. & Schulze, D. J. (1988). Eclogite-facies ultramafic

xenoliths from Colorado Plateau diatreme breccias: comparison with

eclogites in crustal environments, evaluation of the subduction

hypothesis, and implications for eclogite xenoliths from diamondi-

ferous kimberlites. In: Smith, D. C. (ed.) Eclogites and Eclogite-Facies

Rocks. New York: Elsevier, pp. 387–450.

Hoal, K. E. O., Hoal, B. G., Erlank, A. J. & Shimizu, N. (1994).

Metasomatism of the mantle lithosphere recorded by rare

earth elements in garnets. Earth and Planetary Science Letters 126,

303–313.

Humphreys, E., Hessler, E., Dueker, K., Farmer, G. L., Erslev, E. &

Atwater, T. (2003). How Laramide-age hydration of North American

lithosphere by the Farallon slab controlled subsequent activity in the

western United States. International Geology Review 45, 575–595.

Hunter, W. C. (1979). The Garnet Ridge and Red Mesa

kimberlitic diatremes, Colorado Plateau: geology, mineral chemistry,

and geothermobarometry. Ph.D. thesis, The University of Texas at

Austin, 230 pp.

Hunter, W. C. & Smith, D. (1981). Garnet peridotite from

Colorado Plateau ultramafic diatremes: hydrates, carbonates, and

comparative geothermometry. Contributions to Mineralogy and Petrology

76, 312–320.

Hyndman, R. D. & Peacock, S. M. (2003). Serpentinization of the

forearc mantle. Earth and Planetary Science Letters 212, 417–432.

Jackson, S. E., Pearson, N. J., Griffin, W. L. & Belousova, E. (2004).

The application of laser-ablation microprobe inductively coupled

plasma mass-spectrometry (LAM–ICPMS) to in-situ U–Pb zircon

geochronology. Chemical Geology 211, 47–69.

Krogh, E. J. (1988). The garnet–clinopyroxene geothermometer:

a reinterpretation of existing experimental data. Contributions to

Mineralogy and Petrology 99, 44–48.

Krogh Ravna, E. (2000). The garnet–clinopyroxene Fe2þ–Mg

geothermometer: an updated calibration. Journal of Metamorphic

Geology 18, 211–219.

Laird, J. (1988). Chlorites: metamorphic petrology. In: Bailey, S. W.

(ed.) Hydrous Phyllosilicates. Mineralogical Society of America, Reviews in

Mineralogy 19, 405–453.

Leach, T. M. & Rodgers, K. A. (1978). Metasomatism in the Wairere

serpentinite, King Country, New Zealand. Mineralogical Magazine

42, 45–62.

Lee, C.-T., Yin, Q., Rudnick, R. L. & Jacobsen, S. B. (2001).

Preservation of ancient and fertile lithospheric mantle beneath the

southwestern United States. Nature 411, 69–73.

Ludwig, K. R. (2001). Users Manual for Isoplot/Ex rev. 2.49: a geochrono-

logical toolkit for Microsoft Excel. Berkeley Geochronology Center, Special

Publication 1a.

Luth, R. W. (2004). Mantle volatiles—distribution and consequences.

In: Carlson, R. W. (ed.) Treatise on Geochemistry, Volume 2, The Mantle

and Core. New York: Elsevier, pp. 319–361.

McGetchin, T. R. & Silver, L. T. (1970). Compositional relations in

minerals from kimberlite and related rocks in the Moses Rock dike,

San Juan County, Utah. American Mineralogist 55, 1738–1771.

McGetchin, T. R. & Silver, L. T. (1972). A crustal–upper-mantle

model for the Colorado Plateau based on observations of crystalline

rock fragments in the Moses Rock dike. Journal of Geophysical Research

77, 7022–7037.

Mercier, J. C. (1976). Single-pyroxene geothermometry and geobaro-

metry. American Mineralogist 61, 603–615.

Minier, J. & Reiter, M. (1991). Heat flow on the southern Colorado

Plateau. Tectonophysics 200, 51–66.

O’Hanley, D. S., Schandl, E. S. & Wicks, F. J. (1992). The origin of

rodingites from Cassiar, British Columbia, and their use to estimate

T and P (H2O) during serpentinization. Geochimica et Cosmochimica Acta

56, 97–108.

Pearce, J. A. & Peate, D. W. (1995). Tectonic implications of the

composition of volcanic arc magmas. Annual Review of Earth and

Planetary Sciences 23, 251–285.

Philippot, P. & Selverstone, J. (1991). Trace-element-rich brines in

eclogitic veins: implications for fluid composition and transport

1922

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005

during subduction. Contributions to Mineralogy and Petrology 106,

417–430.

Proyer, A., Dachs, E. & McCammon, C. (2004). Pitfalls in geothermo-

barometry of eclogites: Fe3þ and changes in the mineral chemistry of

omphacite at ultrahigh pressures. Contributions to Mineralogy and Petrology

147, 305–318.

Rice, J. M. (1983). Metamorphism of rodingites: Part I. Phase relations

in a portion of the system CaO–MgO–Al2O3–SiO2–CO2–H2O.

American Journal of Science 283-A, 121–150.

Rivalenti, G.,Mazzucchelli,M., Barbieri,M.A., Parenti,M., Schmid,R.

& Zanetti, A. (1997). Garnetite-forming processes in the deep crust:

the Val Fiorina case study (Ivrea–Verbano Zone, NW Alps). European

Journal of Mineralogy 9, 1053–1071.

Roden, M. F. (1981). Origin of coexisting minette and ultramafic

breccia, Navajo volcanic field. Contributions to Mineralogy and Petrology

77, 195–206.

Roden, M. F. & Shimizu, N. (2000). Trace element abundances

in mantle-derived minerals which bear on compositional complex-

ities in the lithosphere of the Colorado Plateau. Chemical Geology

165, 283–305.

Roden, M. F., Smith, D. & Murthy, V. R. (1990). Chemical constraints

on lithosphere composition and evolution beneath the Colorado

Plateau. Journal of Geophysical Research 95, 2811–2831.

Rubatto, D. & Hermann, J. (2003). Zircon formation during fluid

circulation in eclogites (Monviso, Western Alps): implications for Zr

and Hf budget in subduction zones. Geochimica et Cosmochimica Acta

67, 2173–2187.

Schandl, E. S., O’Hanley, D. S. & Wicks, F. J. (1989). Rodingites in

serpentinized ultramafic rocks of the Abitibi greenstone belt, Ontario.

Canadian Mineralogist 27, 579–591.

Scherer, E. E., Cameron, K. L., Johnson, C. M., Beard, B. L.,

Barovich, K. M. & Collerson, K. D. (1997). Lu–Hf geochronology

applied to dating Cenozoic events affecting lower crustal xenoliths

from Kilbourne Hole, New Mexico. Chemical Geology 142, 63–78.

Schulze, D. J., Helmstaedt, H. & Carmichael, D.M. (1987). Hydrated

pyroxenite xenoliths from Navajo diatremes, Colorado Plateau:

pressure–temperature estimates and metasomatic reactions in the

mantle wedge above the Farallon Plate. Geological Society of America,

Abstracts with Programs 19, 835.

Selverstone, J., Pun, A. & Condie, K. C. (1999). Xenolithic evidence

for Proterozoic crustal evolution beneath the Colorado Plateau.

Geological Society of America Bulletin 111, 590–606.

Semken, S. C. &McIntosh, W. C. (1997). 40Ar/39Ar age determinations

for the Carrizo Mountains laccolith, Navajo Nation, Arizona. New

Mexico Geological Society Guidebook, 48th Field Conference, pp. 75–80.

Severinghaus, J. & Atwater, T. (1990). Cenozoic geometry and thermal

state of the subducting slabs beneath western North America.

Geological Society of America, Memoirs 176, 1–22.

Smith, D. (1979). Hydrous minerals and carbonates in peridotite

inclusions from the Green Knobs and Buell Park kimberlitic dia-

tremes on the Colorado Plateau. In: Boyd, F. R. & Meyer, H. O. A.

(eds) The Mantle Sample: Inclusions in Kimberlites and Other Volcanics.

Washington, DC: American Geophysical Union, pp. 345–356.

Smith, D. (1995). Chlorite-rich ultramafic reaction zones in Colorado

Plateau xenoliths: recorders of sub-Moho hydration. Contributions to

Mineralogy and Petrology 121, 185–200.

Smith, D. & Levy, S. (1976). Petrology of Green Knobs diatreme,

New Mexico, and implications for the mantle below the Colorado

Plateau. Earth and Planetary Science Letters 19, 107–125.

Smith, D. & Zientek, M. (1979). Mineral chemistry and zoning in

eclogite inclusions from the Colorado Plateau. Contributions to

Mineralogy and Petrology 69, 119–131.

Smith, D., Connelly, J. N., Manser, K., Moser, D. E., Housh, T. B.,

McDowell, F. W. &Mack, L. E. (2004). Evolution of Navajo eclogites

and hydration of the mantle wedge below the Colorado Plateau,

southwestern United States. Geochemistry, Geophysics, Geosystems 5,

10.1029/2003GC000675.

Spencer, J. E. (1996). Uplift of the Colorado Plateau due to lithosphere

attenuation during Laramide low-angle subduction. Journal of

Geophysical Research 101, 13595–13609.

Sun, S-s. & McDonough, W. F. (1989). Chemical and isotopic syste-

matics of oceanic basalts: implications for mantle composition and

processes. In:Saunders,A.D.&Norry,M. J. (eds)Magmatism in theOcean

Basins. Geological Society, London, Special Publications 42, 313–345.

Switzer, G. S. (1975). Composition of garnet xenocrysts from three

kimberlite pipes in Arizona and New Mexico. Smithsonian Contributions

to the Earth Sciences 19, 1–21.

Thompson, J. B. (1959). Local equilibrium in metasomatic processes.

In: Abelson, P. H. (ed.) Researches in Geochemistry. New York: John

Wiley, pp. 427–457.

Trommsdorff, V. & Evans, B. W. (1969). The stable association

enstatite–forsterite–chlorite in amphibolite facies ultramafics of

the Lepontine Alps. Schweizerische Mineralogische und Petrographische

Mitteilungen 49, 325–332.

Trommsdorff, V. & Evans, B. W. (1974). Alpine metamorphism of

peridotitic rocks. Schweizerische Mineralogische und Petrographische

Mitteilungen 54, 333–352.

Ueda, H., Usuki, T. & Kuramoto, Y. (2004). Intraoceanic unroofing of

eclogite facies rocks in the Omachi Seamount, Izu–Bonin frontal

arc. Geology 32, 849–852.

Usui, T., Nakamura, E., Kobayashi, K., Maruyama, S. &

Helmstaedt, H. (2003). Fate of the subducted Farallon plate inferred

from eclogite xenoliths in the Colorado Plateau. Geology 31, 589–592.

Vrana, S. & Fryda, J. (2003). Ultrahigh-pressure grossular-rich

garnetite from the Moldanubian Zone, Czech Republic. European

Journal of Mineralogy 15, 43–54.

Wang, L., Esssene, E. J. & Zhang, Y. (1999). Mineral inclusions in

pyrope crystals from Garnet Ridge, Arizona, USA: implications for

processes in the upper mantle. Contributions to Mineralogy and Petrology

135, 164–178.

Wendlandt, E., DePaolo, D. J. & Baldridge, W. S. (1993). Nd and Sr

isotope chronostratigraphy of Colorado Plateau lithosphere: implica-

tions for magmatic and tectonic underplating of the continental

crust. Earth and Planetary Science Letters 116, 23–43.

Wendlandt, E., DePaolo, D. J. & Baldridge, W. S. (1996). Thermal

history of Colorado Plateau lithosphere from Sm–Nd mineral

geochronology of xenoliths. Geological Society of America Bulletin 108,

757–767.

West, M., Ni, J., Baldridge, W. S., Wilson, D., Aster, R., Gao, W. &

Grand, S. (2004). Crust and upper mantle shear wave structure of

the southwest United States: implications for rifting and support for

high elevation. Journal of Geophysical Research 109, B03309,

doi:10.1029/2003JB002575.

Wiedenbeck, M., Alle, P., Corfu, F., Griffin, W. L., Meier, M.,

Oberli, F., von Quadt, A., Roddick, J. C. & Spiegel, W. (1995).

Three natural zircon standards for U–Th–Pb, Lu–Hf trace element

and REE analyses. Geostandards Newsletter 19, 1–24.

Wilson, D., Aster, R., Baldridge, W. S., Gao, W., Gok, R., Grand, S.,

Ni, J., Sandvol, E., Semken, S. & West, M. (2003). Imaging

crust and upper mantle seismic structure in the southwestern

United States using teleseismic rceiver functions. Leading Edge 22,

232–237.

Woodhead, J. D., Hergt, J. M., Davidson, J. P. & Eggins, S. M. (2001).

Hafnium isotope evidence for ‘conservative’ element mobility during

1923

SMITH AND GRIFFIN GARNETITE XENOLITHS, COLORADO PLATEAU

subduction zone processes. Earth and Planetary Science Letters 192,

331–346.

Zhang, R. Y., Liou, J. G. & Zheng, J. P. (2004). Ultrahigh-pressure

corundum-rich garnetite in garnet peridotite, Sulu terrane, China.

Contributions to Mineralogy and Petrology 147, 21–31.

Zheng, J., Griffin, W. L., O’Reilly, S. Y., Lu, F., Wang, C., Zhang,

M., Wang, F. & Li, H. (2004a). 3�6Ga lower crust in central

China: new evidence on the assembly of the North China craton.

Geology 32, 229–232.

Zheng, J., Griffin, W. L., O’Reilly, S. Y., Lu, F., Yu, C., Zhang, M. &

Li, H. (2004b). U–Pb and Hf-isotope analysis of zircons in mafic

xenoliths from Fuxian kimberlites: evolution of the lower crust

beneath theNorthChinaCraton.Contributions toMineralogy and Petrology

148, 79–103.

1924

JOURNAL OF PETROLOGY VOLUME 46 NUMBER 9 SEPTEMBER 2005