Journal of Petrology Advance Access originally published online on September 20, 2006
Journal of Petrology 2006 47(12):2463-2489; doi:10.1093/petrology/egl051
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Origin and Source Evolution of the Leucite Hills Lamproites: Evidence from Sr–Nd–Pb–O Isotopic Compositions
1 GEOLOGICAL SURVEY OF CANADA 601 BOOTH STREET, OTTAWA, ONTARIO, CANADA K1A 0E8
2 OTTAWA–CARLETON GEOSCIENCE CENTRE, DEPARTMENT OF EARTH SCIENCES, CARLETON UNIVERSITY OTTAWA, ONTARIO, CANADA K1S 5B6
RECEIVED JUNE 17, 2003; ACCEPTED AUGUST 21, 2006
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONWhole-rock major and trace element and O, Sr, Nd and Pb isotopic data are reported for 3·0–0·89 Ma lamproites from the Leucite Hills, Wyoming, USA. The two main groups of lamproites, madupitic lamproites and phlogopite lamproites, are geochemically distinct and cannot be related to one another by either fractional crystallization or crustal contamination. It seems likely that the geochemical differences between these two rock types are related to variations in source mineralogy and depth of partial melting. The high Mg-number and large ion lithophile element abundances and negative
Nd values of the lamproites indicate a mantle source that has experienced stages of both depletion and enrichment. The negative Nb, Ta and Ti anomalies in mantle-normalized trace element diagrams and low time-integrated U/Pb, Rb/Sr and Sm/Nd ratios of both lamproite groups and other Cenozoic igneous rocks from the Wyoming Archean Province indicate an ancient metasomatic enrichment (>1·0 Ga) of the mantle source associated with the subduction of carbonate-bearing sediments. Other chemical characteristics of the Leucite Hills lamproites, especially their high K2O and volatile contents, are attributed to more recent metasomatism (<100 Ma) involving influx from upwelling mantle during back-arc extension or plume activity. KEY WORDS: isotopes; lamproites; metasomatism; Leucite Hills; Wyoming
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INTRODUCTION |
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Lamproites are Mg-rich igneous rocks (MgO > 5 wt %) that are peralkaline [(K2O + Na2O)/Al2O3 > 1 (molar)], perpotassic [K2O/Al2O3 > 1 (molar)] and ultrapotassic [K2O/Na2O > 3 (molar)]. They span a compositional range of 45–55 wt % SiO2, 4–10 wt % Al2O3, 1–5 wt % TiO2, 2–10 wt % CaO, 5–10 wt % K2O, 0·2–1·5 wt % Na2O and 0·5–2·0 wt % P2O5 (Mitchell & Bergman, 1991
). Although volumetrically insignificant relative to other potassium-rich igneous rocks, lamproites have attracted a great deal of attention among petrologists because of their unusual geochemistry, distinctive mineralogy and potential to contain diamonds. Lamproites occur in intraplate settings (e.g. USA, Greenland, India, Australia, Antarctica) and in rare instances in post-collisional environments (e.g. SE Spain and central Italy). The Leucite Hills, Wyoming, USA, is considered the type locality for lamproites (Mitchell & Bergman, 1991). Even though the Leucite Hills lamproites (LHL) have been extensively studied for more than a century (e.g. Zirkel, 1867; Cross, 1897; Carmichael, 1967; Fraser, 1987; Mitchell, 1995a; Lange et al., 2000), isotope data and trace element concentrations have rarely been reported for the same sample. Considering the distinctive geochemical characteristics of the Leucite Hills lamproites, none of the previous studies have successfully and convincingly explained the nature and the evolutionary history of their mantle sources.
The geochemical characteristics of lamproites, such as high Mg-number and Ni, low Al2O3, and very high concentrations of incompatible elements (higher than those of any other alkaline igneous rocks), are consistent with a mantle source that has undergone multiple depletion and enrichment events (Fraser, 1987). The major problem surrounding the petrogenesis of magmas of lamproitic affinity involves whether their ultimate source is in the lithospheric or sub-lithospheric mantle. Models proposed for the petrogenesis of lamproites range from partial melting of metasomatized lithosphere (e.g. Foley, 1992; Nelson, 1992) to melting of subducted, continentally-derived sediments stored within the Transition Zone at the base of the upper mantle (e.g. Ringwood et al., 1992; Murphy et al., 2002).
In this study, the nature of the mantle source of the LHL is investigated in an attempt to understand the range of chemical (e.g. depletion, enrichment) and geodynamic (e.g. subduction, mantle plume activity) processes that may have contributed to its evolution. This involves a detailed evaluation of the geochemical characteristics of the LHL and integration of data from other Cenozoic volcanic rocks erupted onto the Archean Wyoming craton.
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GEOLOGY AND SAMPLE DESCRIPTION |
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The Leucite Hills are located at 41°47'N, 109°00'W, NE of Rock Springs in Wyoming, USA (Fig. 1). The lamproites in the Leucite Hills cut through the northern flank of the Upper Cretaceous Rock Springs Uplift and clastic sedimentary rocks that belong to the Eocene Fort Union, Wasatch and Green River Formations (Johnston, 1959
; Smithson 1959; Bradley, 1964). All of these rocks overlie a thick cratonic basement known as the Wyoming craton. Composed mainly of supracrustal rocks and granitic gneiss terrains, the Wyoming craton experienced multiple episodes of deformation and metamorphism from 
3·0 to 2·5 Ga (e.g. Stuckless et al., 1985; Hofmann, 1989; Frost et al., 1998). Other tectonic events, including folding and faulting, and magmatism also affected the Wyoming craton during the Proterozoic, resulting in a highly fractured, sheared, anisotropic and heterogeneous lithosphere (Blackstone, 1983). Although many of the basement fractures were reactivated during the Phanerozoic, no major deformation of the basement occurred until the late Cretaceous with the onset of the Laramide Orogeny (Blackstone, 1983).
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The western United States experienced an abrupt increase in volcanic activity in Eocene to Pleistocene times. In the Wyoming craton, magmatism began around 55 Ma (Absaroka volcanism) and continued until 3·0–0·89 Ma with the eruption of the LHL (Mcdowell, 1971
; Mitchell & Bergman, 1991; Lange et al., 2000). The LHL consist of 22 volcanic occurrences that are mostly oriented NW–SE or parallel to structural features such as the Farson Lineament and Maastrichtian thrust fault, which are products of the Laramide Orogeny (Blackstone, 1983; Fig. 1). The lamproites crop out as small groups of volcanic cones, lava flows, plugs and dykes. Highly vesicular scoria and cinders are the dominant materials that make up the tephra cones, whereas flow units and some pyroclastic deposits form composite cones. From field observations, it appears that explosive activity either post-dated or occurred at the same time as the effusive activity. |
PETROGRAPHY AND CLASSIFICATION |
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The petrography of the LHL has been described previously by Cross (1897)
, Kemp & Knight (1903), Kuehner (1980), Carmichael (1967), Fraser (1987) and Mitchell & Bergman (1991). Therefore, only brief petrographic descriptions of the samples used in this study are given here.
The lamproites that were studied consist mainly of a combination of phlogopite, diopside, sanidine, leucite, apatite, perovskite, and minor K-rich richterite, wadeite, priderite and sherbakovite. Of these minerals, phlogopite, clinopyroxene, and apatite can occur as either phenocrysts or as groundmass crystals, whereas the remainder usually crystallize as groundmass phases. Because lamproites from different parts of the Earth exhibit different and distinct mineralogies and textures, a number of diverse names have been used to classify them. To avoid the use of complicated and unnecessary locality names, the studied samples have been classified using the nomenclature proposed by Scott Smith & Skinner (1984a, 1984b), Mitchell (1985), Woolley et al. (1996) and Le Maitre (2002) (Table 1). As a result, the lamproites from the Leucite Hills can be classified into four groups: (1) diopside–leucite-phlogopite lamproites; (2) diopside–sanidine–phlogopite lamproites; (3) madupitic lamproites; (4) transitional madupitic lamproites. The diopside–leucite–phlogopite lamproites and diopside–sanidine–phlogopite lamproites are genetically related, represent the same magma type (Ogden, 1979; Mitchell & Bergman, 1991), and are collectively referred to as phlogopite lamproites. The madupitic lamproites and transitional madupitic lamproites also show many geochemical affinities (Mirnejad, 2002) and the two are grouped under the heading madupitic lamproites.
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ANALYTICAL METHODS |
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The whole-rock major and trace element compositions of the Leucite Hills samples were measured using inductively coupled plasma (ICP) and ICP–mass spectrometry (MS) facilities, respectively, at Activation Laboratories Ltd., Ancaster, Ontario. Samples were taken into solution using standard fusion techniques. A description of the sample preparation can be found at http://www.actlabs.com. Reproducibility, based on repeat analyses, for the major elements is ±1·0% of the quoted values, whereas most trace elements have an uncertainty of ±5·0% if the concentrations are >100 ppm and ±10% for those with concentrations <100 ppm.
For the radiogenic isotopic analyses, whole-rock powders were digested in Savillex beakers using ultrapure HF and HNO3 in the proportion 2:5 and were heated for 24 h to ensure complete digestion. Rb, Sr and rare earth elements (REE) were separated from other elements using ion exchange chromatography (Bell & Simonetti, 1996). The procedure is based on HCl elution through 20 cm Teflon columns of 1·0 cm internal diameter filled with 13 ml of Bio-Rad AG50W-X8 cation exchange resin (200–400 mesh). Pb was separated from other elements using ion exchange chromatography with HCl, HBr and nanopure H2O elution through 0·2 and 0·5 ml columns that were filled with AG1-X8 cation exchange resin (100–200 mesh). The average values of total procedural blanks obtained during the study period are as follows: Rb 0·5 ng, Sr 1·5 ng, Sm 0·04 ng, Nd 0·26 ng and Pb 1 ng.
Sr, Nd and Pb isotopic ratios were determined using a Finnigan-MAT 261 multi-collector mass spectrometer at Carleton University operated in the static mode. Samples were loaded on either Ta or Re filaments. A standard was run with every 12 samples. Based on numerous runs during the course of this study, the following average values were obtained for the standards: NBS-987 87Sr/86Sr = 0·710251 ± 0·00003, La Jolla 143Nd/144Nd = 0·511870 ± 0·00003, NBS-981 206Pb/204Pb = 16·890 ± 0·010, 207Pb/204Pb = 15·429 ± 0·013, and 208Pb/204Pb = 36·498 ± 0·042. Uncertainties are given at the 2
level. An average fractionation correction of 0·12 ± 0·01% per mass unit was applied to all measured Pb isotopic ratios based on analyses of NBS-981. Measured Sr and Nd isotopic ratios for the unspiked samples were corrected for fractionation to an 88Sr/86Sr ratio of 8·3752 and a 146Nd/144Nd ratio of 0·7219. Reproducibility for both Nd and Sr isotopic ratios is ±0·004% of the quoted values.
For O isotopic analysis, 7 mg of the whole-rock sample was oven-dried for 24 h. Samples were then transferred to Ni bombs and fluorinated with BrF5. The Ni bombs were heated at 600°C for 12 h. Extracted O2 was converted to CO2 and then measured on a Finnigan-MAT 252 multi-collector mass spectrometer at the Geological Survey of Canada, Ottawa. The 18O/16O ratios of the samples were normalized to an internal standard and then to V-SMOW. The reproducibility of O isotopic ratios based on multiple runs of NCSU-Qtz is ±0·2
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CHEMICAL COMPOSITION |
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Major and trace element analyses of the LHL samples are given in Tables 2 and 3. Additional major and trace element data can be found in papers by Carmichael (1967)
, Kuehner (1980), Fraser (1987) and Mitchell & Bergman (1991). The major element (wt %) compositions of the LHL based on our new data are: 41·02–55·82 SiO2, 7·20–9·99 Al2O3, 4·13–6·41 Fe2O3T, 2·05–2·73 TiO2, 3·30–12·72 CaO, 4·74–11·54 K2O, 0·54–1·85 Na2O and 1·30–2·99 P2O5. These values are close to the ranges for lamproites cited by Mitchell & Bergman (1991). The Mg-numbers are all very similar and range from 0·72 to 0·79.
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Madupitic lamproites from Pilot Butte and Badgers Teeth have the highest MgO and some of the lowest SiO2 contents of any of the LHL. The highest P2O5 concentrations are shown by the madupitic lamproites from the Badgers Teeth locality. Figure 2 shows the variation of Fe2O3T vs SiO2 (Fig. 2a) and Al2O3 vs CaO (Fig. 2b). In both diagrams the LHL show negative correlations and form two distinct groups with the madupitic lamproites having higher Fe2O3T and CaO and lower SiO2 and Al2O3 contents than the phlogopite lamproites. For comparison, the compositions of lamproites from Western Australia, Antarctica (Gaussberg), Spain and Italy are also plotted in Fig. 2. The Western Australia and the Gaussberg lamproites have higher Fe2O3T contents, whereas most of the lamproites from Italy and Spain have higher Al2O3 contents than the LHL.
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Trace element concentrations in the LHL are given in Table 3. Ba ranges from 4470 to 11 690 ppm, Sr from 1830 to 7233 ppm, Rb from 166 to 296 ppm, and La from 119 to 402 ppm. As with other lamproites, the Ni (104–333 ppm) and Cr (136–560 ppm) abundances in the LHL are greater than in many other alkaline igneous rocks. The average concentrations of Ta and Nb in the madupitic lamproites are 6 and 125 ppm, respectively, whereas for the phlogopite lamproites they are 3 and 47 ppm, respectively. Madupitic lamproites also have higher average Th (45 ppm) and U (10 ppm) contents than the phlogopite lamproites (Th 16 ppm, U 4 ppm). The Sr concentrations of the madupitic lamproites range from 2965 to 7233 ppm, values that are higher than the range found in the phlogopite lamproites (1830–2149 ppm).
One of the characteristic features of the LHL is their high concentration of REE and high light REE/heavy REE (LREE/HREE) ratios [(La/Yb)N = 150–600]. The LREE concentrations and the (La/Yb)N ratios are shown in Fig. 3a; these are higher in the madupitic lamproites than in the phlogopite lamproites (Fig. 3a, inset). A compilation of REE data from lamproites worldwide (Mitchell & Bergman, 1991) shows that the LHL do not differ greatly from other lamproite types in terms of REE abundance and LREE enrichment.
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Trace element data for the LHL normalized to primitive mantle values are plotted on a conventional mantle-normalized trace element diagram (Fig. 3b) in order of increasing compatibility from left to right. In general, the madupitic lamproites are more enriched in the large ion lithophile elements (LILE) and high field strength elements (HFSE) than the phlogopite lamproites. The normalized trace element patterns for the various lamproite types show similar enrichments in many of the incompatible trace elements as well as pronounced Nb, Ta and Ti troughs (Fig. 3b).
The combined trace and major element data show that the two groups of the LHL are chemically distinct from one another. Trace element ratios are also distinctive for the phlogopite and madupitic lamproites, including Ba/La, Ba/Th, Ba/Nb, Rb/Nb, Nb/U, Ce/Pb, and Tb/Yb ratios (not shown). The Ce/Pb ratios in the LHL are variable and range from 7 to 25 (average 14), values that almost extend from near the average value of four for the continental crust to 25 for the mantle (Hofmann, 1997).
The whole-rock Sr, Nd and Pb isotopic compositions of the lamproites from this study are listed in Table 4 and plotted in Figs 4 and 5. The Nd–Sr isotopic compositions of the LHL all lie in the enriched quadrant (Fig. 4) and define two groups, one corresponding to the phlogopite lamproites and the other to the madupitic lamproites. From Table 4 and Fig. 4 it can be seen that the madupitic lamproites have higher 143Nd/144Nd ratios (0·51203–0·51211, Nd –11·7 to –10·1) than the phlogopite lamproites (0·51186–0·51194, Nd –15·0 to –13·5). 87Sr/86Sr ranges from 0·70534–0·70563 for the madupitic lamproites to 0·70566–0·70606 for the phlogopite lamproites. The LHL form a vertical array at relatively constant 87Sr/86Sr and variable 143Nd/144Nd ratios. Although the LHL have similar Nd values to lamproites from Gaussberg, Spain, Italy and West Kimberley, their 87Sr/86Sr ratios are lower.
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The madupitic lamproites have Pb isotopic ratios (206Pb/204Pb 17·4–17·58, 207Pb/204Pb 15·47–15·50, 208Pb/204Pb 37·42–37·52) somewhat higher than those of the phlogopite lamproites (206Pb/204Pb 17·13–17·23, 207Pb/204Pb 15·46–15·48, 208Pb/204Pb 37·20–37·32) (Fig. 5). Unlike the Pb data from the LHL, those from Western Australia, Gaussberg, Spain and Italy plot above the Stacey–Kramers growth line on the 207Pb/204Pb vs 206Pb/204Pb diagram (Fig. 5a). Data for all of these lamproites plot above the Northern Hemisphere Reference Line (NHRL).
Some new oxygen isotopic data from the LHL are given in Table 4. The 
18OSMOW of the samples fall within a relatively narrow range from +8·21 to +8·90, and no significant differences are observed between the madupitic lamproites and phlogopite lamproites from different locations.
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THE RELATIONSHIP BETWEEN PHLOGOPITE LAMPROITES AND MADUPITIC LAMPROITES |
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The observation that the LHL form two distinct geochemical groups in most of the major and trace element and radiogenic isotope diagrams means that the madupitic and phlogopite lamproites cannot simply be related by fractional crystallization. Alternative explanations might involve crustal assimilation, variation in depth of partial melting and/or source mineralogy.
Crustal contamination
Although crustal contamination has a minimal effect on the chemical composition of lamproites because of their extreme enrichment in incompatible elements as well as their high Ni and Cr contents (Carmichael, 1967; Kuehner, 1980; Vollmer et al., 1984), Ogden (1979) suggested that the madupitic lamproites from Pilot Butte might have attained their distinctive mineralogy and chemical composition by reaction between a phlogopite lamproite melt and entrained crustal rocks, a hypothesis solely based on petrographic and field observations. We tested this using a simple binary mixing model.
Contamination can occur in two ways, one involving mixing between crustal- and mantle-derived melts, and the other by assimilation of crustal rocks. Here, the bulk assimilation model is evaluated using simple binary mixing between upper crustal rocks and lamproite melts. Although a more realistic approach involves the assimilation–fractional crystallization (AFC) model as formulated by DePaolo (1981), the overall outcome is very model dependent, involving such parameters as the composition of the initial magma, the composition of the host rock, the composition of the fractionating phases, the partition coefficients between the fractionating phases and the residual melt, and the rate of assimilation and crystallization (DePaolo, 1981). Because of the lack of constraints on many of these parameters only a bulk assimilation model is considered here.
Wedepohl's (1995) estimate of the composition of the upper continental crust was used as an end-member in the binary mixing calculations. Figure 6a shows that bulk assimilation by the phlogopite lamproites of upper crust increases the Al2O3, MnO, and Na2O abundances and decreases the TiO2, MgO, K2O and P2O5 contents of the contaminated magma. In this case, the major element composition of the contaminated phlogopite lamproite magma fails to approach the chemical composition of any of the existing madupitic lamproite types in the Leucite Hills for up to 50% contamination (f = 0·5, where f is the weight fraction and is calculated as f = A/(A + B); A and B are the weight fractions of the contaminant and the original magma, respectively). Similarly, bulk assimilation of upper crust by the madupitic lamproites cannot generate lamproite melts whose major element contents resemble those of the phlogopite lamproites.
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The trace element concentrations of the contaminated magmas, normalized to the average trace element concentration of the phlogopite lamproites, are plotted in Fig. 6b. Contamination by upper crust decreases the concentration of most incompatible elements and increases the concentration of the HREE in the melt. The incompatible element characteristics of the resulting contaminated phlogopite lamproite melt differ significantly from those of the madupitic lamproites. A similar approach shows that contamination of a madupitic lamproite melt by upper crust cannot generate magmas with trace element contents similar to those of the phlogopite lamproites.
The Sr–Nd isotopic data place additional constraints on crustal assimilation models. Contamination by upper crust (Fig. 7a) increases the 87Sr/86Sr ratios of the madupitic lamproites to values comparable with those of the phlogopite lamproites but only for f > 0·5 (i.e. assimilation of >50% of crustal material). The Mg-number and Ni contents of the madupitic lamproites are much too high to allow for such degrees of assimilation. Similarly, the madupitic lamproite melt needs to assimilate substantial amounts of upper crustal rocks (f > 0·5) for its 143Nd/144Nd ratio to reach values close to of those of the phlogopite lamproites (0·5118) (Fig. 7b). However, even though contamination of the madupitic lamproite could generate a melt with a Nd isotopic composition similar to that of the phlogopite lamproites, the major and trace element composition of such a melt are still very different. Assimilation of upper crust by a melt with the composition of a phlogopite lamproite cannot generate a melt with Nd isotopic ratios similar to the madupitic lamproite (Fig. 7b).
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On the basis of these simple mixing models, it is unlikely that the two major lamproite types found in the Leucite Hills (i.e. phlogopite lamproite and madupitic lamproite) are related to one another, either by fractional crystallization or by crustal contamination. It is more likely that the geochemical distinctions of these two lamproites represent source variations.
Source mineralogy and depth of partial melting
One explanation for variations in the chemical compositions of mantle-derived melts could be the melting of different mineral phases contained within metasomatic veins trapped within the lithosphere (e.g. Meen et al., 1989; Foley, 1992). The results of melting reactions for the cratonic vein assemblage phlogopite + clinopyroxene + amphibole (Foley et al., 1999) show that amphibole melts completely within 50°C of the solidus and thus the melt compositions are primarily controlled by the composition of the amphibole. Such melts are expected to be richer in SiO2 and K2O and poorer in CaO and Al2O3 than melts derived from peridotites at similar pressures. Foley et al., however, noted that the consumption of accessory mineral phases such as apatite, ilmenite and rutile can strongly influence melt compositions. The resulting melt from the vein assemblage can infiltrate and react with the surrounding peridotite wall-rock. Mitchell & Edgar, (2002) attributed the difference in chemical composition between the phlogopite lamproites from North Table Mountain and madupitic lamproites from Middle Table Mountain to different degrees of partial melting of vein plus wall-rock assemblages.
Geochemical variations in the two types of lamproites from Leucite Hills can also be related to depth of partial melting. The anhydrous kalsilite–forsterite–quartz system has been applied to the origin of lamproites (Kuehner, 1980; Kuehner et al., 1981) to show that madupitic lamproites are generated by lower degrees of partial melting and at greater depths than the phlogopite lamproites. This system, however, has some limitations because it does not contain any Ca-bearing and hydrous phases.
Melting experiments on mantle materials (e.g. Hirose & Kushiro, 1993; Baker & Stolper, 1994; Wasylenki et al. 1996; Herzberg et al., 2000) indicate that high-pressure mantle-derived melts normally have higher Fe2O3T but lower SiO2 and Al2O3 contents than those generated at lower pressures. An increase in Fe2O3T and Tb/Yb and a decrease in Al2O3 and SiO2 of basalts from west to east across the Basin and Range province were thought to indicate shallower melting beneath the western than the central parts (Wang et al., 2002). From the data presented in Fig. 2 it can be seen that the madupitic lamproites have higher Fe2O3T and lower SiO2 and Al2O3 contents than the phlogopite lamproites. If these compositional variations can be related to results obtained from the melting experiments of Hirose & Kushiro (1993), Baker & Stolper (1994), Wasylenki et al. (1996) and Herzberg et al. (2000), then the madupitic lamproite melts were generated at greater depths relative to the phlogopite lamproite.
Variations in melting depths are also supported by the trace element composition, 40Ar/39Ar dating, and volumetric proportions of the two lamproite types from the Leucite Hills. Recent high-pressure experimental data show that with increasing pressure, clinopyroxene becomes richer in the Ca-Tschermaks component and behaves similarly to garnet in differentiating the REE (Salters & Longhi, 1999). Because melts in equilibrium with garnet have high Tb/Yb values as a result of garnet's stronger preference for HREE and middle REE (MREE) relative to spinel, magmas with high Tb/Yb ratios should be derived from deeper parts of the mantle. From the data presented in Table 3, madupitic lamproites are characterized by marginally higher (Tb/Yb)N ratios (13·8) than the phlogopite lamproites (10·4). The higher LREE concentration in the madupitic lamproites than in the phlogopite lamproites provides further evidence for a deeper origin for the madupitic lamproites. In addition, the oldest lavas in the Leucite Hills, the madupitic lamproites (3·0–2·5 Ma old) form <1% of the volume of lamproites while the younger phlogopite lamproites (1·8–0·8 Ma old) form >99% (Lange et al., 2000), observations that may indicate that the madupitic lamproites represent the first batch of partial melt generated at greater depths and at lower degrees of partial melting of the mantle source. Although this model might explain the chemical differences between the phlogopite and madupitic lamproites from Leucite Hills, the very different isotopic compositions for the two groups requires that the vein minerals at lithospheric depths must have been out of isotopic equilibrium with the surrounding mantle. Variations in the source mineralogy associated with the depth of partial melting can clearly explain the isotopic compositions of the madupitic and phlogopite lamproites as well as the trends of the isotopic data observed in Figs 9 and 10. Throughout their evolutionary history the two lamproitic melts did not interact with one another, and hence were able to preserve their distinct chemical identities.
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LHL IN RELATION TO OTHER CENOZOIC IGNEOUS ROCKS FROM THE WYOMING CRATON |
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Comparisons between the LHL and other igneous rocks of Eocene to Paleocene age from the Wyoming craton (Fig. 1) are shown in Figs 8–10. These rocks form a number of different magmatic suites, all <55 Ma in age. The Absaroka volcanic province of northwestern Wyoming and southwestern Montana consists of two sub-parallel, NW-trending belts: the western Washburn (andesite, basaltic andesite and dacite) and the eastern Sunlight (shoshonite to latite) volcanic rocks (55–52 and 50–48 Ma, respectively) (Feeley et al., 2002
; Feeley & Cosca, 2003). The Independence suite of volcanic rocks, which crops out within the eastern Absaroka volcanic field, is 52–49 Ma old and contains low-K high-alumina tholeiitic basaltic andesite to shoshonite and to high-K dacite (HATS series), and a high-magnesium andesite to banakite (HAB series) (Meen & Eggler, 1987; Harlan et al., 1996). On the northern margin of the Wyoming craton, the Missouri Breaks diatreme (50–45 Ma) contains kimberlite and alnöite (Scambos, 1987), and the Crazy Mountains complex (48 Ma) consists of alkalic malignite, phonolite and trachyte, and subalkalic basalt, andesite and rhyolite (Dudás et al., 1987; Dudás, 1991). The volcanic rocks of the Bearpaw Mountains (54–50 Ma), located in northern Montana, are dominated by minettes and latites (Macdonald et al., 1992), whereas in central Montana, the Highwood Mountains magmatic province (53–50 Ma) consists of quartz-normative latites overlain by flows and intruded by stocks and dikes of potassic mafic rocks (O'Brien et al., 1991). The 27 Ma Smoky Butte lamproite (Marvin et al., 1980), located NE of Leucite Hills, consists of armalcolite–phlogopite and sanidine–phlogopite lamproites (Fraser, 1987; Mitchell et al., 1987). About 200 km NW of the Leucite Hills lies the Yellowstone volcanic province, whose activity can be traced over a period of 19 Myr, from southwestern Idaho to the present-day locus of the Yellowstone Plateau (Smith & Braile, 1994). The Yellowstone Plateau and the Snake River Plain volcanic fields, two major products of the Yellowstone hotspot, consist of subalkaline basalt–rhyolite associations (Hildreth et al., 1991).
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In a primitive mantle-normalized trace element diagram (Fig. 8) most of the Cenozoic volcanic rocks show Nb, Ta and Ti depletions, except the Missouri Breaks rocks. In terms of their Sr and Nd isotopic compositions, almost all of the Tertiary–Quaternary volcanic rocks from Wyoming plot in the enriched quadrant (excluding the data from Missouri Breaks) and define a near-vertical trend (Fig. 9), with the Yellowstone basalts lying at one end and the Smoky Butte lamproites at the other. Similar to the data from the LHL, those from the Independence volcanic field show large variations in
Nd and define two distinct groupings (Fig. 9). It is interesting to note that all of the Cenozoic volcanic rocks from the Wyoming craton have moderately radiogenic Sr isotopic compositions, indicating that the source region of these magmas had time-integrated Rb/Sr slightly higher than that of bulk Earth (BE). It is also clear from Fig. 9 that the isotopic signatures of the alkaline rocks from the Wyoming Province, and hence their sources, are distinct from the K-rich volcanic rocks of Spain, Italy, Gaussberg, and Western Australia.
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On a 207Pb/204Pb vs 206Pb/204Pb diagram the Wyoming Cenozoic igneous rocks plot to the left of the geochron and below the Stacey–Kramers growth line (Fig. 10a). These rocks also plot above the NHRL and form a linear trend on 207Pb/204Pb vs 206Pb/204Pb and 208Pb/204Pb vs 206Pb/204Pb diagrams (Fig. 10a and b). The most and least radiogenic Pb isotopic signatures of any of the Tertiary–Quaternary igneous rocks emplaced into the Wyoming craton are from the Missouri Breaks kimberlite–alnöite suite and the Smoky Butte lamproites, respectively. In general, the isotopic variations among the Cenozoic igneous rocks from the Wyoming craton point to a widespread heterogeneous mantle source. No Pb isotope data are available at present for the Absaroka and Bearpaw volcanic rocks.
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PETROGENESIS OF THE LHL |
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Any model proposed for the origin of the LHL has to take into account the following:
- higher abundances of MgO, Ni and Cr, and lower abundances of CaO, Al2O3 and Na2O in lamproites relative to most primitive, mantle-derived melts;
- the two distinct chemical groupings of the madupitic and phlogopite lamproites;
- the LREE enrichment and strongly negative Nd signature of the LHL and some of the other Cenozoic igneous rocks from the Wyoming craton;
- the moderately radiogenic Sr isotope compositions of the LHL as well as other Cenozoic igneous rocks from the Wyoming craton despite wide variations in K2O contents and Rb/Sr ratios;
- the Pb isotopic compositions of the LHL and other contemporaneous igneous rocks from the Wyoming craton that plot to the left of the geochron and below the Stacey–Kramers growth line;
- the presence of both small- and large-scale source heterogeneity, reflected by the variations in the Sr–Nd–Pb isotopic compositions of the LHL and other Cenozoic volcanic rocks found throughout the Wyoming craton;
- the Nb, Ta and Ti depletions and LILE enrichments of the LHL as well as other Cenozoic volcanic rocks from the Wyoming craton.
On the basis of these observations, we propose a three-stage model (Fig. 11) involving: (1) major melt removal from the Wyoming upper mantle in the Archean (2·8 Ga) that left behind a refractory and depleted sub-continental mantle lithosphere; (2) ancient (>1 Ga) metasomatism of the sub-continental mantle related to subduction of crustal materials; (3) recent (<100 Ma) metasomatism of the sub-continental lithosphere during mantle upwelling leading to K enrichment and subsequent melting of the metasomatized portions of the mantle source.
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Depletion processes
The low abundances of Al2O3, SiO2, and CaO and the high MgO, Ni and Cr contents of the studied samples (Tables 2 and 3) are consistent with the geochemical characteristics of other lamproites (e.g. Bergman, 1987
; Foley, 1992), supporting derivation of the LHL from a major-element depleted mantle source, dominantly harzburgitic in composition. In addition, the relatively low HFSE and HREE contents of the LHL are common among lamproites and a variety of other alkaline igneous rocks from Wyoming and are generally thought to indicate mantle source depletion by previous partial melting events (Eggler et al., 1988; O'Brien et al., 1995). Although the low concentrations (see Table 3) of Tm (<0·2 ppm), Yb (<0·9 ppm) and Lu (<0·11 ppm) in the LHL probably indicate residual garnet in the source, calculations by Tainton & McKenzie (1994) show that to account for these concentrations, the mantle source of the lamproites must have been depleted by extensive melting.
The involvement of a refractory (harzburgitic) mantle source for lamproites is supported by data from mantle xenoliths and phase equilibria studies. Although mantle xenoliths are exceptionally rare in lamproites, those so far recovered are dominantly dunite and garnet harzburgite (Atkinson et al., 1984; Eggler et al., 1987; Mitchell & Bergman, 1991; Carlson & Irving, 1994). Refractory mantle xenoliths, dominantly harzburgites, have been reported from the Colorado–Wyoming kimberlites (Eggler et al., 1987), and from the Williams kimberlite, associated with a more extensive Eocene suite of alkalic ultramafic intrusions in the Missouri Breaks; these indicate that much of the lithospheric mantle is depleted in magmatophile elements (i.e. Ca, Al and Fe) (Carlson et al., 1999). In addition, ultramafic xenoliths from the Bearpaw Mountains are composed mainly of harzburgites, and rare dunites and werhlites, with major element compositions moderately to extremely depleted relative to the chemical composition of average peridotite from the sub-continental lithosphere (Downes et al., 2004). Most of the published phase equilibrium data for natural lamproites (e.g. Foley, 1993; Mitchell, 1995a; Edgar & Mitchell, 1997; Mitchell & Edgar, 2002), as well as studies of anhydrous and volatile-bearing (H2O, CO2, F) compositions in the synthetic system kalsilite–forsterite–quartz at high pressures also point to harzburgitic protoliths as the mantle source of ultrapotassic melts (Kuehner, 1980; Foley et al., 1986).
Generation of a refractory harzburgitic mantle beneath the Wyoming craton can be attributed to the formation of continental crust at between 3·5 and 2·5 Ga, as indicated by the ages of the Granite Mountain gneisses (Wooden & Mueller, 1988) and the Wind River Range granites (Stuckless et al., 1985; Frost et al., 1998). Based on geothermobarometry calculations on mantle xenoliths from Wyoming–Colorado kimberlites, Eggler et al. (1987) proposed that the entire Wyoming sub-continental mantle to the depth of at least 200 km was depleted during a melting event or events in the Precambrian. Those workers argued that although the Front Range granitic batholiths have been dated at 1·4 and 1·0 Ga (Peterman et al., 1968), they largely reflect remobilization of pre-existing crust. This, coupled with the lack of widespread basaltic volcanism during the Proterozoic, might suggest that the depletion event was restricted to the Archean. Additional evidence for Archean crustal extraction comes from mafic crustal xenoliths from the Leucite Hills. Based on geochemical compositions and whole-rock Pb–Pb and Sm–Nd pseudoisochrons (Mirnejad & Bell, in preparation) and U–Pb dating of zircons (Farmer et al., 2005) indicating ages of about 2·8 Ga, these rocks are interpreted to represent fragments of igneous material intruded into the deep crust and thus support the Archean depletion of the Wyoming upper mantle. Other support for Archean depletion is given by Os isotopic ratios of spinel peridotite, pyroxenite and glimmerite xenoliths in Eocene minette dikes from southern Montana that yield model ages of 2·7 to 2·9 Ga (Carlson & Irving, 1994; Rudnick et al., 1999). Although mantle xenoliths from the Bearpaw Mountains have provided somewhat younger (2·45–1·13 Ga) Nd model ages (Carlson & Irving, 1994; Downes et al., 2004), these might represent a mixed age intermediate between that of the Archean depletion event and a younger LREE enrichment event.
Enrichment processes
The mantle source of the LHL clearly experienced enrichment as witnessed by the high alkali metal and LREE contents (Table 3), extremely negative Nd values, and high Ba/Nb, Ba/La, Nb/Pb and Ce/Pb ratios.
Lamproites, along with kimberlites, are the most extreme products of mantle enrichment processes (Hawkesworth et al., 1985) and of the many hypotheses proposed to produce the high incompatible element and volatile contents in their mantle source, metasomatism is the most widely accepted (e.g. Menzies & Wass, 1983; Hawkesworth et al., 1985; Foley et al., 1986; Menzies, 1987; Menzies & Hawkesworth, 1987). The addition of hydrous melts or fluids enriched in K2O and incompatible elements to the mantle source is a prerequisite for the generation of lamproitic melts (Hawkesworth et al., 1990). Based on melting experiments of primary lamproites and the near-solidus phase relationships at high P–T conditions (4–7 GPa, 1000–1200°C; Foley, 1992), phlogopite ± richterite ± clinopyroxene ± apatite ± titanate are considered to be the main minerals making up the metasomatic veins in the lamproite mantle source (Foley, 1993; Mitchell, 1995a; Edgar & Mitchell, 1997; Mitchell & Edgar, 2002). In addition, the stability of hydrous phases under upper mantle conditions has been experimentally investigated by a number of workers (e.g. Kushiro et al., 1967; Tronnes et al., 1988; Konzett et al., 1997; Sato et al., 1997), demonstrating that phlogopite and amphibole can be stable throughout the cratonic lithosphere. Carlson & Irving (1994) and Rudnick et al. (1999) reported glimmerite veins composed of mica, apatite, orthopyroxene, clinopyroxene, rutile, zircon, monazite, magnetite, and chromite that cut harzburgitic xenoliths contained in potassic volcanic rocks in the Highwood Mountains. In addition, Carlson & Irving (1994) have indicated that most peridotite xenoliths from the Highwood Mountains contain mica and amphibole. It is generally agreed that vein assemblages formed during metasomatism of the mantle contribute greatly to the major and trace element budgets of potassic or ultrapotassic melts (Luth et al., 1993; Schmidt et al., 1999).
Enrichment processes leading to the formation of metasomatic mantle minerals are diverse and can be caused by the reaction of mantle rocks with metasomatic fluids or melts derived from the dehydration or partial melting of subducted slabs (Hawkesworth et al., 1990; Murphy et al., 2002), or volatiles or melts that emanate from mantle upwellings (McKenzie, 1989; Tainton & McKenzie, 1994). On the basis of our findings we propose that both an early metasomatism of the lithospheric mantle source related to subduction processes and a contribution from a recent mantle upwelling or plume were involved in the enrichment of the mantle below the Wyoming craton.
Ancient metasomatic event
An ancient metasomatic event in the mantle source is indicated by the extremely negative Nd values of the LHL (Fig. 9). With regard to the approximate timing of this enrichment event, both late Archean and mid-Proterozoic have been cited as the age of metasomatism. Vollmer et al. (1984) argued that metasomatism took place during stabilization of the Archean lithosphere in the Wyoming Province at about 2·7 Ga and showed that slight enrichment in Nd/Sm and Rb/Sr ratios in the upper mantle was sufficient to generate the observed Nd and Sr isotopic variations seen in the LHL. Even the Pb isotopic data for the Leucite Hills and Smoky Butte lamproites are consistent with ancient metasomatism; these data plot to the left of geochron and below the Stacey–Kramers growth line (Fig. 10a) indicating U/Pb and Th/Pb fractionation of their sources. Based on the isotopic compositions of the Smoky Butte lamproites, Fraser (1987) and Mitchell et al. (1987) suggested that depletion of U, Th and Pb in the source occurred subsequent to a 2·5 Ga metasomatic event that enriched the Wyoming mantle in LREE, U, Th and Rb. Archean enrichment of the Wyoming lithosphere has also been documented from potassic mafic magmas of the Highwood Mountains (O'Brien et al., 1995), involving an increase in Ba/Nb and Ba/La, and decrease in Nb/Pb and Ce/Pb ratios along with a decrease in Nd values from –11 to –20. Not all of the evidence, however, supports Archean enrichment. The Os, Sr, Nd and Pb isotopic compositions of mica harzburgite xenoliths, and zircon and monazite dates from glimmerite veins from the Highwood Mountains give ages of about 1·8 Ga, indicating a mid-Proterozoic enrichment event (Carlson & Irving, 1994; Rudnick et al., 1999). Some additional Rb–Sr model ages for phlogopite from the Bearpaw Mountains mantle xenoliths yield values of about 1·25 Ga (Downes et al., 2004), and model ages for mafic alkalic and subalkalic rocks from the Crazy Mountains fall between 1·3 and 1·8 Ga (Dudás et al., 1987; Dudás, 1991). Metasomatic Nd model ages of 2·2–2·0 Ga (Feeley, 2003) were obtained from mafic lavas from Sunlight and Washburn volcanoes. Eggler et al. (1988), on the basis of model Nd ages (TDM) of a variety of Tertiary igneous rocks from the Wyoming craton, came to a similar conclusion. Websterite and pyroxenite mantle xenoliths from the Wyoming–Colorado kimberlite pipes were interpreted by Eggler et al. (1987) as representing metasomatic zones created during an enrichment event, during either the Archean or the Proterozoic (1·6–1·7 Ga). The Nd values for the madupitic and phlogopite lamproites from the Leucite Hills are much too low to have been generated from a non-metasomatized, harzburgitic source, suggesting that the Sm/Nd ratio must have been lowered by metasomatism. Model ages calculated for the LHL samples from this study, assuming a CHUR reservoir, correspond to 715 Ma for the madupitic lamproites and 1034 Ma for the phlogopite lamproites. These ‘ages’ represent minimum limiting values, and give support to an old metasomatism of the mantle source associated with LREE enrichment. Given the different estimates of the timing of enrichment of the Wyoming lithosphere, and the possibility of more than one period of metasomatism, we refer to this >1 Ga episode as the ancient metasomatic event.
The similar trace element patterns of the LHL and other Tertiary volcanic rocks emplaced within the Wyoming craton (excluding Missouri Breaks) (Fig. 8) suggest that the enrichment of their mantle sources may be related. The Nb, Ta and Ti depletions and the LILE and LREE enrichment of the LHL closely resemble those observed in convergent margin tectonic settings (Pearce, 1982; Cox, 1988; Hawkesworth et al., 1990), indicating that the metasomatic signature in their mantle source was probably subduction-related. In addition, the average Ce/Pb ratio (17) of the LHL is much lower than that of mid-ocean ridge basalts (MORB) and ocean island basalts (OIB) (25), supporting the involvement of fluids or melts from subducted crustal materials.
Although lamproites are rarely found in convergent tectonic settings and few can be linked to modern subduction zones, the patterns of source enrichment indicated by the trace element and Sr–Nd–Pb isotopic compositions of the LHL nevertheless would suggest subduction-related processes (Fig. 8). In this context, Bergman (1987) noted that lamproites are commonly located above fossil Benioff Zones and Mitchell & Bergman (1991) suggested that the tectonic settings and compositional characteristics of lamproites such as Nb and Ta depletion and LILE and LREE enrichment point to the involvement of ancient subducted slabs. Other evidence from Tertiary igneous rocks in the Wyoming Province supports subduction-related mantle metasomatism, including xenocrystic plagioclase and granitic melt inclusions in zircons from glimmerite veins that cut harzburgite xenoliths, as well as crust-like REE patterns, and Nb depletions of xenoliths from the Highwood Mountains (Rudnick et al., 1999). Both Carlson & Irving (1994) and Rudnick et al. (1999) used zircon and monazite ages of about 1·8 Ga from these glimmerite veins to date the involvement of an old subduction process. An increase in Ba/Nb, Ba/La, Nb/Pb and Ce/Pb ratios and a decrease in Nd values among the Highwood volcanic rocks were considered by O'Brien et al. (1995) as evidence that the ancient metasomatic signatures in their mantle source are subduction-related.
It is clear from the Sr–Nd isotopic diagram (Fig. 9) that the LHL plot in the enriched quadrant, and in the case of Pb (Fig. 10) to the left of the geochron and below the Stacey–Kramers growth line. The observed Pb isotopic variation can be attributed to the extent to which U/Pb ratios are fractionated during subduction. It is possible that high O2 fugacity during subduction oxidized U4+ to the more soluble U6+ and caused U to be lost relative to Pb, thus lowering the U/Pb ratio in a downgoing slab (White & Dupré, 1986). The Pb, Sr and Nd isotopic signatures of the LHL can also be attributed to the chemical heterogeneity of subducted materials, such as crustally derived sediments containing carbonate and phosphate. Because of their low Rb/Sr, U/Pb ratios, and relatively high Nd/Sm ratios (e.g. Othman et al., 1989; Plank & Langmuir, 1998) any fluids or melts derived from such sediments can metasomatize the lithosphere and generate the low time-integrated Nd, Sr and Pb isotopic ratios found in the Leucite Hills lamproites.
There is also field evidence for late Archean or mid-Proterozoic subduction beneath the Wyoming and adjacent cratons. Based on geochemical data and geological observations, Mueller & Wooden (1988), Wooden & Mueller (1988), Frost et al. (1998) and Chamberlain et al. (2003) have suggested that the Wyoming craton may have been the site of a long-lived, late Archean, convergent continental margin. Evidence for Archean subduction includes compositionally diverse, late Archean andesitic amphibolites in the Beartooth Mountains in Wyoming (Mueller et al., 1983) with trace element patterns that closely resemble those of modern arc magmas. A number of late Archean supracrustal sequences (mafic volcanic rocks, metagraywackes, pelitic schists, and minor intermediate to felsic volcanic rocks and quartzites) preserved within the southern and western portions of the Wyoming Province are consistent with intracratonic and cratonic margin settings for basin development that occurred during crustal growth along the SW margin of the Wyoming Province (Chamberlain et al., 2003). As an alternative, Bennet & De Paolo (1987), Hoffman (1989), and Bickford et al. (1990), have suggested that the western United States experienced collisional tectonics and associated magmatism during the mid-Proterozoic (1·85–1·65 Ga). U–Pb zircon geochronology of metasomatized mantle xenoliths from the Great Falls tectonic zone in the NW margin of the Wyoming craton led Carlson & Irving (1994), Carlson et al. (1999) and Rudnick et al. (1999) to propose a model in which metasomatism of the Archean lithospheric mantle of the Wyoming craton was associated with the creation of the Great Falls tectonic zone, accompanied by accretion of mobile belts to its northern boundary around 1·8 Ga.
In summary, subduction of crustal materials (including sediments) during either the late Archean or mid-Proterozoic seems to be the most likely process responsible for the metasomatism of the lithospheric keel beneath the Archean Wyoming craton prior to 1·0 Ga.
Recent metasomatic event
In addition to Archean depletion and ancient (>1 Ga) enrichment events, there is also sufficient evidence for the involvement of recent (<100 Ma) metasomatic activity affecting the Wyoming mantle.
Most lamproites worldwide have radiogenic Sr isotopic compositions indicating high, long time-integrated Rb/Sr ratios in their source (Fig. 4). McCulloch et al. (1983) ascribed the highly radiogenic Sr isotopic compositions of lamproites from Western Australia to Proterozoic (>1 Ga) Rb enrichment relative to Sr. The 40Ca/42Ca ratios of the Sisimiut lamproites, West Greenland (Nelson & McCulloch, 1989), indicate that K2O enrichment within their mantle source occurred at least 1 Gyr prior to lamproite eruption. In contrast, however, the 87Sr/86Sr ratios of the LHL are only moderately radiogenic and this may indicate that the ancient metasomatism produced a mantle source characterized by low Rb/Sr ratios. Carlson & Irving (1994) considered mantle xenoliths with high negative Nd values and moderately radiogenic Sr isotopic ratios from the Highwood Mountains as evidence for ancient metasomatism characterized by low Rb/Sr and Nd/Sm ratios. Such geochemical characteristics, common among all the Cenozoic igneous rocks from the Wyoming craton (Fig. 10), including those of Smoky Butte and the Absaroka volcanic province, have been attributed by Mitchell et al. (1987) and Feeley (2003) to insignificant amounts of Rb being added to the mantle source during the ancient metasomatic event.
The relatively low 87Sr/86Sr ratios and thus low time-integrated Rb/Sr ratios in the source of the LHL are seemingly at odds with the relatively high K2O and Rb contents of the lamproites. Examples of enrichment in K2O without accompanying Sr isotopic enrichment are also indicated by the highly K2O-enriched lavas from the Sunda arc of Indonesia (Wheller et al., 1987). The marked differences between the 87Sr/86Sr ratios in the West Kimberley and Smoky Butte lamproites (see Fig. 9) was considered by Mitchell et al. (1987) to reflect different styles of metasomatism, with the Smoky Butte source containing more richterite and titanite and less phlogopite than the West Kimberley source. Mitchell et al. (1987) also emphasized that alumina-deficient and K2O-poor richterites are unable to provide sufficient Al2O3 and K2O to account for the abundant phlogopite and sanidine contained in the Smoky Butte lamproites, and suggested recent introduction of phlogopite into the mantle source of the Smoky Butte lamproites during volatile fluxing in the Tertiary as a more viable explanation for the different styles of enrichment between the Smoky Butte and West Kimberley lamproites. A similar argument can thus be used to explain the differences in Sr isotope compositions between the Leucite Hills and West Kimberley lamproites.
The whole-rock oxygen isotope data show that the 18O values of the LHL (18O = +8 to +9 Table 4) are higher than mantle values (18O = +5·5 Mattey et al., 1994). Our 18O data are very similar to the values of +8·8 reported by Kuehner (1980) for phlogopite phenocrysts from the LHL. Because high-level, crustal contamination is ruled out in the petrogenesis of the LHL, the 18O of the phlogopite phenocrysts may be considered a source feature and probably reflects recent metasomatic activity associated with hydrous fluids or melts. Further evidence for the recent addition of K2O to the source of the LHL may come from the lack of correlation between the K2O contents and Nd isotopic ratios of the studied samples. As with the LHL, the Washburn and Sunlight mafic rocks from the Absaroka volcanic province show wide variations in K2O, LILE, LREE contents and Rb/Sr ratios with no correlation with 143Nd/144Nd ratios, features that are consistent with recent (<100 Ma) mantle metasomatism (Feeley, 2003). As can be seen in Fig. 12, there is no correlation between Nd values and K2O contents of the Cenozoic igneous rocks from the Wyoming craton. This is strong evidence that the K2O enrichment of the Wyoming mantle source was independent of the ancient LREE enrichment, and that the addition of K2O probably occurred during a much younger metasomatic event. The question as to why Sr–Nd–Pb isotopes were not affected by the recent metasomatic activity remains to be answered, but it could be that the K-bearing fluid phase had very low Nd, Sr and Pb contents.
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Other evidence for recent metasomatism in the region includes the presence of undeformed phlogopite in mantle xenoliths from the Bearpaw Mountains (Downes et al., 2004
) that was introduced after tectonism and before incorporation of the xenoliths into their host magmas, perhaps during the early Tertiary. Also, Carlson et al. (1999) have noted that the Sr, Nd and Pb isotopic compositions of clinopyroxenes in some of the peridotite xenoliths from the Williams kimberlite overlap those of alkalic magmas from the Montana high-K igneous province, and concluded that these pyroxenes were introduced by metasomatism shortly before the capture and transport of the xenoliths. Those workers also attributed small veins of phlogopite in some xenoliths to recent metasomatic activity. Although there is abundant evidence to support the existence of recent metasomatic activity that led to the addition of K and volatiles to the mantle source of the LHL, the ultimate source of the potassium, and the reasons for preservation of the older isotopic signatures need to be answered. These issues are discussed in the following sections.
Lithospheric vs sub-lithospheric sources
Isotopic and trace element variations among the Cenozoic igneous rocks from the Wyoming craton (excluding the Missouri Breaks) fall outside the known ranges of OIB and therefore their parental melts cannot be derived solely from a normal, sub-lithospheric mantle source. Moreover, for metasomatic minerals to retain their geochemical integrity from the time of the ancient metasomatic event they need to be preserved in a reservoir that has been isolated from the convecting mantle for a long period of time. Such a reservoir could be located in the rigid sub-continental lithosphere, within the mantle Transition Zone or even the lower mantle.
In those models for lamproite petrogenesis involving the Transition Zone, subducted continentally derived sediment, stored for long periods of time at the base of the upper mantle (650 km), can undergo partial melting and provide melts capable of metasomatizing the mantle. Ringwood et al.'s (1992) model for the origin of kimberlites and lamproites involved subducted sediments that formed a garnetite layer within the Transition Zone. Generation of partial melts within the Transition Zone as a result of convection currents from the lower mantle resulted in an LREE-enriched melt capable of metasomatizing the overlying depleted mantle. Ringwood et al. (1992), based on experiments using a synthetic Group I kimberlite, showed that at 16 GPa and 1650°C, corresponding to P–T conditions within the Transitions Zone, majorite garnet (13% Al2O3) and ß-M2SiO4 were the liquidus or near-liquidus phases. Majorite garnet inclusions in diamond (Moore & Gurney, 1991) and exsolution features in garnet–clinopyroxene xenoliths from the Jagersfontein kimberlites (Haggerty & Sautter, 1990) provided additional evidence that the garnet originally contained a majorite component. Ringwood et al. (1992) argued that the Transition-Zone model could apply to lamproite petrogenesis because of the geochemical similarity between olivine lamproites and Group II kimberlites.
Although there is now abundant seismic tomographic evidence for the accumulation of subducted oceanic slabs at the 650 km discontinuity (e.g. Simons et al., 1999; Fukao et al., 2001), the assumption that lamproites can originate from such sources should be treated with caution. The Jagersfontein kimberlites from which majorite garnets were recovered are Group I kimberlites; these do not share the distinctive geochemical characteristics of lamproites and Group II kimberlites (Smith, 1983; Mitchell, 1989, 1995b; Skinner, 1989). Moreover, olivine is dominantly a xenocryst phase in olivine lamproites and thus the compositions of olivine lamproites do not reflect those of their parental magmas (Mitchell, 1995a). In addition, Tainton & McKenzie (1994) argued that because the majorite garnet exsolution rate constants are unknown, estimates of the time taken for majorite exsolution from garnet cannot be made. Therefore, the majorite garnet inclusions are, at present, only evidence that diamonds, rather than kimberlites, form in the Transition Zone (Tainton & McKenzie, 1994). Although Ringwood et al. (1992) reported high-pressure, near-liquidus garnet and clinopyroxene phases, the chemical compositions of these minerals differ from those of the eclogites or garnetite found as xenoliths in kimberlites (Edgar & Mitchell, 1997). Many high P–T melting experiments on lamproites do not confirm a garnetite source but point to the presence of phlogopite, K-rich amphibole and apatite (Foley, 1992; Mitchell, 1995a; Edgar & Mitchell, 1997; Mitchell & Edgar, 2002) that are necessary to account for the extreme chemical compositions of lamproites and their volatile contents. Experiments on the stability of phlogopite and K-amphibole under upper mantle conditions (e.g. Kushiro et al., 1967; Konzett et al., 1997; Sato et al., 1997) show that phlogopite and amphibole can be stable in the sub-cratonic mantle only to depths of about 250 km.
The model proposed by Murphy et al. (2002) to explain the petrogenesis of the Gaussberg lamproites in the East Antarctic shield is similar, in many ways, to that of Ringwood et al. (1992), involving the melting of subducted, Archean continent-derived sediments stored within the Transition Zone or lower mantle for 2–3 Gyr. In this model, the main potassium-bearing phase at such depths is K-hollandite. Murphy et al. (2002) argued that the low Nd values of the Gaussberg lamproites require long-term source enrichment in the LREE, and the Pb isotope compositions that plot to the left of the geochron indicate U/Pb fractionation during Archean subduction of sediments into the Transition Zone. One of the more important objections made by Murphy et al. (2002) to metasomatism of the Gaussberg sub-continental lithosphere is the lack of any geochemical evidence for recent enrichment. This is not true for the LHL because of the evidence for the recent metasomatism in their mantle source.
The lack of widespread ultrapotassic magmatism in oceanic environments and the confinement of kimberlites and lamproites to mantle sources associated with Archean cratons (White et al., 1995; Jaques & Milligan, 2004) indicate that the origin of lamproites and kimberlites is controlled by thickened lithosphere. Phase equilibria studies suggest that lamproites originate from partial melting of veined mantle sources at 4–7 GPa and 1000–1200°C (Foley, 1993; Mitchell, 1995a; Edgar & Mitchell, 1997; Mitchell & Edgar, 2002), P–T conditions that are very different from those in the Transition Zone. Moreover, the fact that the isotopic compositions of some xenoliths from the sub-continental mantle lithosphere overlap with those of lamproites might be consistent with sub-continental mantle lithosphere as the ultimate source of lamproites (Wilson, 1989).
Geodynamic setting
Recent metasomatism and subsequent partial melting of the Wyoming mantle can be explained by (1) shallow subduction of the Farallon plate under the Wyoming sub-continental lithosphere, or (2) upwelling of mantle materials during lithospheric extension. These are discussed in turn.
The traditional interpretation of Cenozoic magmatism in the western USA is that active subduction of the Farallon plate led to the generation of a series of mantle-derived magmas (e.g. Lipman, 1980; Bird, 1984). However, more recent models involve mixing between partial melts of ancient metasomatized lithosphere and asthenospheric melts that migrated upward as the result of slab roll-back or slab break-off associated with flat subduction of the Farallon plate (Madsen et al., 2006). For example, the arc-like geochemical characteristics of the Tertiary–Quaternary volcanic rocks of the Wyoming craton, including those from the Elkhead Mountains (Leat et al., 1988), the Highwood Mountains (O'Brien et al., 1991, 1995) and the Absaroka volcanic fields (Feeley, 2003; Feeley & Cosca, 2003) are attributed to metasomatism of the overlying asthenosphere during flat subduction of the Farallon plate in the late Cretaceous.
The existence of the Farallon plate beneath the Montana–Wyoming area and thus enrichment of the mantle wedge during Farallon plate subduction is questionable. To relate the Tertiary–Quaternary magmatism in the Wyoming craton to Farallon plate subduction, the subduction zone must have extended at least 1200 km inland from the continental margin and the angle of subduction would have to have been unusually shallow (e.g. Severinghaus & Atwater, 1990; Lee, 2005). Although Murphy et al. (2003) and Ihinger et al. (2004) considered that overriding of a mantle plume by the Pacific oceanic plate at 50 Ma resulted in shallowing of the subduction zone, the problem with such a model is the distance between the easternmost limit of Eocene magmatism and the inferred location of the trench (1500 km to the west), which exceeds the maximum width of the Andean cordillera associated with flat slab subduction. In addition, the thick and coherent lithospheric mantle beneath the Wyoming craton would have stood as a physical impediment to flat-slab subduction (Dudás, 1991). Heat-flow patterns (Blackwell, 1991), geothermobarometric calculations from xenoliths (Eggler et al., 1987), and seismic tomography data (Artemieva & Mooney, 2001, and references therein) support the presence of thick continental lithosphere beneath the Wyoming craton. Even if the Farallon slab was able to subduct underneath the Wyoming craton, it would have remained in contact with colder lithosphere, probably precluding partial melting. It is more likely that the dominant arc-like geochemical characteristics of the Tertiary volcanic rocks from Wyoming, including those of the LHL, are inherited from an ancient metasomatic component (>1 Ga) in the lithospheric mantle. Fitton et al. (1988) and Hergt et al. (1991) have demonstrated that Nb, Ta and Ti depletions as well as high LILE/HFSE and LREE/HFSE ratios in continental igneous rocks not associated with contemporary subduction systems may preserve a record of the effects of older subduction in the sub-continental lithosphere.
An alternative to models that attribute recent metasomatism to flat subduction of the Farallon plate is recent metasomatism and partial melting of the sub-continental lithosphere by a mantle upwelling or a plume. The Yellowstone hotspot is the only possible candidate for mantle plume activity in Wyoming (e.g. Smith & Braille, 1994; Camp, 1995; Schutt et al., 1998; Smith & Siegel, 2000). Mitchell & Bergman (1991) have stressed the close relationship between the Leucite Hills magmatism and Yellowstone hotspot activity, and suggested that the onset of partial melting in the mantle source of the LHL coincided with the time when the outer parts of the Yellowstone hotspot track passed by the Rock Springs region at 1–2 Ma. Edgar (1983) investigated the relationship between the K/(K + Na) (molecular ratio) and the Ba content of ultrapotassic rocks from the western USA with distance from the Yellowstone hotspot and noticed that the LHL situated closest to Yellowstone have the highest K/(K + Na) and Ba contents. In the model of Edgar (1983), K was introduced from the Yellowstone plume and transported by fluids into the upper mantle.
A difficulty with the hotspot or plume model is the lack of temporal and spatial evidence for the presence of the Yellowstone plume beneath the various volcanic centres in Wyoming. If the Yellowstone plume induced partial melting in the mantle source that generated all of the Tertiary–Quaternary magmatism in Wyoming, then the plume must have existed beneath the Wyoming lithosphere for the last 55 Myr. However, the oldest age proposed for the Yellowstone plume is 19 Ma, an age that marks the time when it emerged near the Oregon–Nevada border (Smith & Braile, 1994; Zoback et al., 1994), much further NW of the present Yellowstone hotspot. Plate reconstructions by Murphy et al. (2003) suggest that at 50 Ma the Yellowstone plume was probably beneath the Pacific ocean floor, close to the continental margin of western North America. It seems unlikely, but not impossible, that at 50 Ma such a plume, located thousands of kilometers to the west of the Bearpaw, Highwood and Smoky Butte localities, could have played a role in the enrichment and partial melting of the mantle source underlying these volcanic fields. Another major drawback to the involvement of a mantle plume are the recent arguments that question the fueling of the Yellowstone hotspot by a plume of hot material rising from the lower mantle (Walker et al., 2004; Anderson, 2005). According to Christiansen (1993), the relations between rift zones in the surrounding areas and the propagation of the Yellowstone hotspot from 19 Ma to the present time appear to be inconsistent with the geometry of a deep mantle plume. In addition, seismic tomography data have not revealed vertical structures with low velocity extending into the lower mantle but a low-velocity body beneath Yellowstone that appears to be restricted to the upper mantle (Christiansen et al., 2002). Anderson (2005), using the polling approach of Courtillot et al. (2003), also concluded that the plume hypothesis for the Yellowstone hotspot scores poorly against an asthenospheric feature associated with stress release induced magmatism.
As an alternative model, the thermal anomaly beneath the thick Wyoming lithosphere could have been created by mantle upwelling independent of plume activity. As favoured by Dudás (1991) for the Crazy Mountains volcanic rocks, the upwelling of mantle below the Wyoming craton could be the result of back-arc extension and lithospheric thinning (decompression) related to Farallon plate subduction (Eggler et al., 1988; Christiansen et al., 2002). Contemporaneous with lithospheric extension, episodic mantle upwelling resulted in partial melting of the heterogeneous Wyoming sub-continental lithosphere that generated most of the Cenozoic magmatism. Beneath Archean cratons, convective upwelling in the underlying sub-lithospheric mantle has been inferred to occur on average every few million years (Foley et al., 1999); thus several partial melting events could have occurred in the mantle. The lithospheric mantle is too cold to melt and perturbation of the mantle solidus by a heat source and/or volatile influx is required before partial melting takes place. A mantle upwelling can spread beneath the lithosphere and raise the temperature to 100–200°C above normal by conduction (White & McKenzie, 1989). However, this is a slow process and more efficient transport of heat as well as volatiles or melts occurs when the upwelling mantle moves upwards through lithospheric channels and fractures (Ebinger & Sleep, 1998), in which case the upwelling mantle viscosity needs to be relatively low (Albers & Christensen, 2001; Xue & Allen, 2005).
Many of the problems associated with the origin of the LHL and their mantle source are analogous to those associated with the petrogenesis of ultrapotassic rocks in Italy; both rock types show depletions in Nb, Ta and Ti, enrichment in the LREE, high Mg-numbers and decoupling between isotopic compositions and some major and trace element abundances. A common feature of both groups of rocks is the involvement of at least two isotopically distinct sources. Continuing debate centres around whether Cenozoic magmatism in Italy is related to subduction or to large-scale plume activity [see discussions by Peccerillo & Lustrino (2005) and Bell et al. (2006) and references therein], similar to the types of models proposed for the origin of the LHL.
The distinct differences in the chemical compositions of the madupitic lamproites and the phlogopite lamproites from the Leucite Hills probably result from variations in the source mineralogy and the depth of partial melting in the metasomatized lithosphere. As discussed previously, the relative volumes of the two groups of lamproites and their temporal and geochemical relationships indicate that the madupitic lamproite magma was generated prior to and at greater depths than the phlogopitic lamproite melts. The heat and volatiles recently introduced by transported materials from the upwelling mantle initially induced metasomatism and partial melting (madupitic lamproites) in the deeper parts of the sub-continental lithosphere. By propagating to shallower levels, the upwelling material underwent further partial melting that resulted in the phlogopite lamproite magmas.
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CONCLUSIONS |
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The two major types of lamproite in the Leucite Hills, madupitic and phlogopite lamproites, show distinct characteristics in many major and trace element and Sr–Nd–Pb isotope diagrams. Although fractional crystallization or crustal contamination has had a minimal effect, variations in the source mineralogy and the depth of partial melting appear to have played an important role in the petrogenesis of the lamproites.
The mantle source involved in generating the LHL has clearly undergone both depletion and enrichment events. The depletion event is related to widespread Archean crustal formation that left behind a refractory, harzburgitic residue. The high Mg-number and Ni and low Al2O3 and Na2O in the lamproites relative to many other alkaline rocks, as well as the results of high P–T phase equilibrium studies are all consistent with the involvement of a harzburgitic mantle source in the genesis of the LHL. However, the high concentration of LILE and LREE in these lamproites indicates that the depletion event was followed by pervasive metasomatism. The negative Nb, Ti and Ta anomalies observed in the LHL trace element patterns indicate the involvement of subducted materials in metasomatizing the mantle source. The ancient age of this subduction-related metasomatism is supported by the Nd isotopic composition of the LHL (i.e. very low Nd values), and evidence from other igneous rocks emplaced in the Wyoming craton also indicates an ancient (>1 Ga) subduction event. As a result of the nature of the subducted materials and/or the high P–T conditions associated with the ancient subduction, the metasomatized mantle source developed moderate Rb/Sr, low U/Pb, and very low Sm/Nd time-integrated ratios. The enriched lithospheric mantle source subsequently behaved as a closed system until recent times (<100 Ma) when it was metasomatized by volatile-rich components derived from mantle upwelling, possibly related to the Yellowstone hotspot or plume, or more probably to back-arc extension associated with subduction of the Farallon plate. The newly added metasomatic components not only provided additional K2O, volatiles and heat to the Wyoming sub-continental lithosphere but also lowered the solidus temperature to initiate partial melting in the source region.
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ACKNOWLEDGEMENTS |
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We are grateful to A. Rukhlov, S. Farrell, D. Gold and E.A. Spencer, and especially B. Kjarsgaard, for constructive discussions and comments. I. de Jong is thanked for her assistance in the clean lab. Help by Bruce Taylor in obtaining oxygen isotopic data is much appreciated. Thanks go to M. Wilson, and the four reviewers, S. Bergman, T. Furman, J. Meen, and D. Prelevic, whose comments added greatly to this manuscript. This work was funded by an NSERC operating grant to K.B.
*Corresponding author. Telephone: (613) 992-4046. Fax: (613) 943-1286. E-mail: hmirneja{at}nrcan.gc.ca
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