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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.
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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
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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.
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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
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405060 Mg
Ca
Fe
50
040
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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.
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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).
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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
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ght p
erce
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Distance (µm)
TiO2
Cr2O3
MnO
(b)
4
8
12
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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?’.
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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
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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
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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.
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