Oxygen isotope, cathodoluminescence, and titanium in quartz geothermometry in the Alta Stock, UT: geochemical insights into pluton assembly and early cooling history

The Alta Stock in the Wasatch Mountains of Utah is a well-studied intrusion and is associated with a large, well-developed contact aureole; thus, the stock provides an excellent opportunity to test new geochemical approaches for investigating the assembly history of plutons. A combination of oxygen isotope analysis, cathodoluminescence imaging, and Ti-in-quartz geothermometry (TitaniQ) was undertaken to document the geochemical signatures of the stock and to examine how these are related to pluton emplacement, growth, and early cooling history. These results suggest a complex assembly and cooling history for the Alta stock. The delta18 O values from whole rock, quartz, feldspar, biotite, magnetite, and hornblende reveal that these minerals have exchanged oxygen isotopes into the subsolidus ( 5 5 0 - 6 5 0 ° C ) . Quartz delta18 0 values range from 8.9%o to 10.0%o, and do not correlate in any obvious way with location within the stock, nor do they correlate with rock type (i.e., granodiorite, mafic enclave, aplite, etc.). Most (0.9%o) of this l.l%o variation may be accounted for by variations in abundance of quartz, the temperature at which it equilibrated with other minerals, and in the extent of subsolidus exchange with hydrothermal fluid. Thus clearly identifying magma increments is difficult in the Alta Stock using major mineral delta18 O values.

OXYGEN ISOTOPE, CATHODOLUMINESCENCE, AND TITANIUM IN QUARTZ GEOTHERMOMETRY IN THE ALTA STOCK, UT: GEOCHEMICAL INSIGHTS INTO PLUTON ASSEMBLY AND EARLY COOLING HISTORY by Benjamin W. Johnson A thesis submitted to the faculty of The University of Utah in partial fulfillment of the requirements for the degree of Master of Science in Geology Department of Geology and Geophysics The University of Utah August 2009 EARL Y III Copyright © Benjamin William Johnson 2009 All Rights Reserved THE U N I V E R S I T Y OF UTAH G R A D U A T E SCHOOL SUPERVISORY COMMITTEE APPROVAL of a thesis submitted by Benjamin William Johnson This thesis has been read by each member of the following supervisory committee and by majority vote has been found to be satisfactory. Chair: John R. Bowman John M. Bartley Barbara P. Nash UNIVERSITY GRADUATE SCHOOL THE U N I V E R S I T Y OF UTAH GRADUATE SCHOOL APPROVAL To the Graduate Council of the University of Utah: I have read the thesis of Benjamin W. Johnson in it s f i n a i fo rm and have found that (1) its format, citations, and bibliographic style are consistent and acceptable; (2) its illustrative materials including figures, tables, and charts are in place; and (3) the final manuscript is satisfactory to the supervisory committee and is ready for submission to The Graduate School. 'Date/ John R. Bowman Chair: Supervisory Committee Approved for the Major Department Chair/Dean Approved for the Graduate Council f Charles A. Wight • Dean of The Graduate School UNIVERSITY SCHOOL FINAL READING APPROVAL ofthe I have read the thesis of Benjamin W. Johnson in its final form ~ /,g~ S!l~~a~ Marjorie A. Chan ~~<. c ~ __ ~J\ • ABSTRACT The Alta Stock in the Wasatch Mountains of Utah is a well-studied intrusion and is associated with a large, well-developed contact aureole; thus, the stock provides an excellent opportunity to test new geochemical approaches for investigating the assembly history of plutons. A combination of oxygen isotope analysis, cathodoluminescence imaging, and Ti-in-quartz geothermometry (TitaniQ) was undertaken to document the geochemical signatures of the stock and to examine how these are related to pluton emplacement, growth, and early cooling history. These results suggest a complex assembly and cooling history for the Alta stock. The 5 O values from whole rock, quartz, feldspar, biotite, magnetite, and hornblende reveal that these minerals have exchanged oxygen isotopes into the sub-solidus 5 5 0 - 6 5 0 ° C ) . Quartz S 1 8 0 values range from o to o, and do not correlate in any obvious way with location within the stock, nor do they correlate with rock type (i.e., granodiorite, mafic enclave, aplite, etc.). Most o) of this l.l%o variation may be accounted for by variations in abundance of quartz, the temperature at which it equilibrated with other minerals, and in the extent of subsolidus exchange with hydrothermal fluid. Thus clearly identifying magma increments is difficult in the Alta 1 o Stock using major mineral 5 O values. OISO sub­solidus ( 550-650°C). OISO 8.9%0 10.0%0, (0.9%0) 1.1%0 Stock using major mineral OISO values. Cathodoluminescence (CL) imaging of quartz grains reveals complex patterns. Clear core-to-rim zoning (consistent with growth zoning) is present only in quartz phenocrysts from the central phase and in samples of the border phase adjacent to the central phase. A bright CL band is present on the outer edge of grains that show clear core-to-rim zoning. Most grains exhibit complex networks of irregular, bright CL bands (interpreted as filled fractures) within dark CL domains (tile-like patterns), mixtures of irregularly-shaped bright-and-dark CL domains (mottled patterns), and less often, enigmatic, irregular zoning. These patterns suggest multiple periods of quartz fracturing and growth, resulting from either synkinematic emplacement of the stock and/or fracturing, cementation, and growth during subsolidus (>600°C) hydrothermal alteration of the quartz. TitaniQ temperatures suggest that core-to-rim zoning in CL records crystal growth during cooling. Temperatures are 800-900°C in the cores of phenocrysts and decrease to 685-700°C in the darker CL zones nearer the exterior of grains. Nearly all phenocrysts have a bright CL rim; this feature is interpreted as a thermal rejuvenation event (Ti temperatures ~800°C). These thermal rejuvenation rims are present through most of the central phase of the Alta stock, and either record a large thermal input into the entire central phase or a series of local heating events induced by the emplacement of a series of magma increments. v fracturing (~600°C) phenocrysts have a bright CL rim; this feature is interpreted as a thermal rejuvenation event (Ti temperatures ~800°C). These thermal rejuvenation rims are present through most of the central phase of the Alta stock, and either record a large thermal input into the entire central phase or a series of local heating events induced by the emplacement of a series of magma increments. v TABLE OF CONTENTS ABSTRACT iv LIST OF TABLES viii LIST OF FIGURES ix ACKNOWLEDGMENTS xi Chapter 1. MAGMA, ISOTOPES, AND CATHODOLUMINESCENCE 1 Introduction ". 1 Oxygen Isotopes 4 TitaniQ/Cathodoluminescence 5 Geologic Setting of the Alta Stock and Emplacement Models 7 2. METHODS 12 Sample Collection, Preparation, and Analytical Procedure 12 3. ANALYTICAL RESULTS 18 Stable Oxygen Isotopes 18 TitaniQ Geothermometry 42 4. DISCUSSION 56 Implications from the Oxygen Isotopic Compositions of the Alta Stock and its Constituent Minerals 56 Igneous Textures Revealed by Cathodoluminescence 65 TitaniQ Temperatures and Early History of the Alta Stock 69 5. CONCLUSIONS 72 Oxygen Isotopes 72 Cathodoluminescence/TitaniQ Temperatures 72 . ...... . ..... . .. . .. . ... .. .. . ... .......... ..... .. .. .. . .. . .... . ... ................ .. ...... .................................................... . ... . ....... .. .. ... ..... ...... ............................................................. . ... ... ... .... ... . ....... . ................................ . ... ..... ..................... ..... ..................... . .. . .. . ...... . . . .. . .. . ....... .. . .. ........ .. .. . ...... .. ..... .. ...... .. ........ .. .... .... .... . .. . .. . ... ..... .... .. . .. ...... .... .. . .. . ... .. ... ...... . .................. TitaniQICathodoluminescence . ... ...... . .... . ... ... ..... .... ... ....... .. . . .... .. .......... ........................ .. .. . ................................................... .. .. ............... . ... ... .... ... ....................... .. .. . .. . .... . ... ...... . .. ... ..... . ............. .. ............ . ..... .... . ...... ..... ...... . ... .. .. .... . .. . .. . ... .. .................... Cathodoluminescence Imaging .. .. ... ... . .. . .... . .. .... . ......... .. ...... . ....... . ...... 35 .. ...... . .. . ...... . .. . . .... ......... . .... . .. ............ . ... ..... .42 ...................................... .. .. ............ . ..................... . .. .. .................. . ....... . ............................. . .. .... . . ...... ................. .. .......... .... .......................... . ........ .... ... ...... . ......................... . ... ......... . ... .. .. .... . .. .. ......... . ................... .. ..... ......... . ... .. .. ..... . .... . .. .......... Titani Q ..... .. ... ... . .. . .... ........... .. . ........ Appendices A: THIN SECTION PHOTOS BY SAMPLE SITE 75 B: OXYGEN ISOTOPE EQULIBRIUM EXCHANGE FACTORS 77 REFERENCES 79 vii ........................................... ........................ .................... .. .............. . ... ..... ........ ............................. ... Vl1 LIST OF TABLES Table Page 3.1. Oxygen isotope data listed by rock type 19 3.2. Calculated oxygen isotope equilibrium temperatures 27 3.3. Average oxygen isotope values 30 3.4. Measured Ti concentration and calculated TitaniQ temperatures 44 4.1. Calculated changes in 6 1 8 0 of minerals as a function of varying Qtz abundance 58 4.2. QEMSCAN mineral modes 59 ............................................ ... .... ..... .. .. . ... ..... ..... .................... . ....... . .. . .... . .. . ...... . ....... ...... .... ..... .44 () 180 ............. .. ..................... .. ........ ..... ........... ..... ........ ... .. .......... . .................. . .. .. .. .. .. ... .... . ......... LIST OF FIGURES gure Page 1.1. Location and simplified geologic map of Alta stock 8 2.1. Outcrop photos of sample sites H2 & H3 13 2.2. Photos of sample site D3 14 2.3. Example of a banded feature at site H7 15 3.1. Oxygen isotope data for whole rock and mineral separates 21 3.2. Plot of Qtz-Bt vs. Qtz-Mag mineral fractionations 22 3.3. Plot of Fsp-Bt vs. Qtz-Bt mineral fractionations 23 3.4. Plot of Qtz-Mag vs. Qtz-Fsp mineral fractionations 24 3.5. Plot of Qtz-Bt vs. Qtz-Fsp mineral fractionations 25 3.6. Qtz or Fsp isotope values vs. whole rock isotope values 28 18 3.7. Outline of the Alta stock showing average WRand Qtzo1 00 data from granodiorite samples 31 3.8. Sample site D3 showing isotope zones and 6 1 8 0 values of quartz and whole rock 33 3.9. Photograph of sample D2B3 34 3.10. CL images of quartz grains from pluton-wide transect 36 3.11. CL images from border phase granodiorite and aplite samples 37 3.12. CL images from sample H3E2 38 3.13. Three pairs of CL and back scattered electron images 40 3.14. CL images of Qtz from 3 zones across the granodiorite-mafic band Figure .. ... ... ... . .. ...... .... . .. . ... ... ... . ...... . .. . .. . .. .. . ... .. .. . ...... .. .. . ... .. .. .... . .. ...... ... ... . . .... ....... .. . ... .. ..... .... 2.3 . .. .... . ... ... . .... .... ... ... . .. . . . ... ........ . . .. . . ......... .. ..... .. .. .. .... . ....... .. ... . .. . .. .. .... . ........ . ... ............. .. .... ... .. ... .. ....... .. . .. . . . ... ... ... . ............... . ............... ............. . ............ 3.7. Outline of the Alta stock showing average WR and Qtz 0180 data from granodiorite samples ........................... .. .. ....... . ... .. .. .... . .... 31 0180 ........... .. ... . .. . .. ... ... ... ... . .......................................... .............. . ...... . .... . ... .... .. . . .... . .. ..... .... .. .. . ... . .. . ... .. .. . ...... . .. . ... ... ... . . .. . .... . .... . .......... . ...................... . ......... .... ... ....... .. ........ . .40 contact ........ .. .. .. . .... ... ... .... .. . ........................... . ....... . ...... .. ...... 41 3.15. CL image with TitaniQ temperatures for a grain from sample D2B3a 49 3.16. Representative CL images from three samples along the pluton-wide transect 50 3.17. CL images with microprobe spots and TitaniQ temperatures from sample H3E2 51 3.18. Outline of the Alta stock with granodiorite sample sites showing average TitaniQ temperatures 53 3.19. TitaniQ temperature data and a representative CL image from site D3 54 3.20. Photograph of sample D2B3 55 4.1. Plot of measured 6 1 8 0 values of Qtz vs. measured amount of Qtz 61 4.2. Measured 6 , 8 0 values of Qtz plotted as a function of the measured oxygen isotope fractionation between Qtz and Bt 62 4.3. EDS spectra taken from a quartz grain in sample 89-1-11 66 ..... .49 ......... .. ........... . ............ ... ............ ... .... .. ....... ........................................................................... ....................................................... ..... ....................................................... 018.. ... .. .. .... 0180 ofQtz . ..... ................ . ......... 11. ... . .. ........... ..... x ACKNOWLEDGMENTS This work would not have been possible save for the help of many. For funding, I thank GSA, PRF, and NSF. I would like to thank Jared Singer, Keith Christianson, Neil Lareau, Krysia Skorko, Will Gallin, and Melinda Hilber for their assistance with the finicky paleomagnetic drill. Many thanks are also due to Quintin Sahratian for help in finding equipment and sample preparation. Eric Thomas and Mike DePhanger also deserve a nod for their help in preparing samples for isotope analysis and microprobe work. I would like to thank John Valley and Mike Spicuzza at the University of Wisconsin for assistance with oxygen isotope analysis. For microprobe work, I must thank Henny Cathey and Mike Spilde at the University of New Mexico. Thanks to Eric Peterson for assistance with the QEMSCAN. My thesis committee has provided invaluable help both in the field and in the lab, and I am extremely grateful for their aid. Finally, I would like to thank my friends and family (Dad, Mom, Tom, and Emily) who provided love and support throughout the entire process. Skol! CHAPTER 1 MAGMA, ISOTOPES, AND CATHODOLUMINESCENCE Introduction Based on outcrop patterns, early geochemical work and modeling, the emplacement of large, granitic plutons has been assumed to occur via diapiric ascent (e.g., Akaad, 1956). Melt forms at depth, rises through the crust, stalls, and solidifies as a single unit. There is a significant amount of speculation when investigating this process, since it is impossible to see it happen. We have not, as of yet, been able to conclusively observe a modern pluton being emplaced. Many of the assumptions associated with this traditional "big tank" emplacement model are now being questioned (e.g., Brown & McClelland, 2000). When a eutectic magma first melts, it is not heated much above its melting temperature. It likely does not have the excess heat required to "melt" its way through the crust to an emplacement level (Glazner, 2007). In addition, stoping of crustal blocks through the magma as a means of transport is not viable at a large scale (Glazner and Bartley, 2006). Glazner and Bartley conclude that a high volume of stoped blocks will freeze a rising magma body, and stoped blocks are likely only minor features within a pluton. Additional latent heat from crystallization could, theoretically, melt the surrounding crustal rock, not only making space for the rising magma but assimilating the crust into the magma (Taylor, 1980). CATHODOLUMINESCENCE 2 Oxygen and strontium isotope data provide evidence for crustal contributions to plutons in a number of magmatic arcs associated with convergent plate margins (Taylor, 1986). However, Taylor (1980) showed that assimilation coupled with fractional crystallization (AFC) is not likely to have formed the observed oxygen and strontium isotope trends of these plutons, at least in the upper crust at or near the level of emplacement. The isotope data show that the magma primarily reflects its source, and not any subsequent process associated with ascent or emplacement in the upper crust (Taylor, 1986). Space making at shallow crustal levels is also a significant issue in pluton formation (Bartley et al., 2008). Field evidence for traditional space-making mechanisms (stoping, wall rock flow, ballooning) is perhaps more tenuous than previously stated (Glazner et al., 2004). It is possible that tectonic extension is the primary space-making mechanism (Hutton, 1982). Typical rates of tectonic extension are ~l-2 mm/yr (Bartley et al., 2008). It is physically impossible at these rates for a large body of magma to accumulate and exist at once; accumulation of a body the size of the Alta stock would require on the order of 106 yrs. In light of these issues, an alternative model of incremental pluton emplacement has been proposed (Pitcher and Berger, 1972; Brown & McClelland, 2000; Mahan et al., 2003; Glazner et al., 2004). This model views large plutons as amalgams of many smaller batches of magma. There are many lines of evidence that favor incremental emplacement. Recent field mapping in plutons has identified the presence of subtle internal contacts that may not have been considered to be of intrusive character previously (Bartley et a l , 2008). Mapped patterns of magnetic susceptibility (a proxy for mineralogic makeup) also suggest a mosaic assembly style (Bartley et al., 2008; Stearns & Bartley, 2008). High- aI., aI., ~ 1-aI., 106 aI., aI., aI., aI., precision U-Pb zircon geochronology has provided a means to constrain directly the time scales of pluton assembly. Recent application of high precision U-Pb zircon geochronology to the Sierra Nevada batholith suggests a protracted history of emplacement of the Tuolumne Intrusive Suite, upwards of several millions of years (Coleman et al., 2004, 2005). Finally, thermal modeling (e.g., Annen et al., 2006a) suggests that pluton emplacement can involve a protracted history defined by periods of emplacement followed by periods of little or no magmatic activity. A significant challenge in validating the incremental emplacement model in the field is that mapping and measuring individual increments is difficult. The prolonged history and many periods of heating and reheating from new batches of magma tend to smear out and/or erase any sharp internal contacts within the intrusion (Bartley et al., 2008). The very process of incremental emplacement erases physical evidence of the process. There are several geochemical/imaging techniques that might prove fruitful in evaluating mechanisms (single-stage vs. incremental) of pluton emplacement, including oxygen isotopes, the newly calibrated TitaniQ (Ti-in-quartz) geothermometer (Wark and Watson, 2006), and cathodoluminescence (CL) imaging of quartz. It is possible that these geochemical tracers and underutilized imaging technique record processes of pluton emplacement that are not easily visible or readily apparent from traditional mapping or thin section petrographic techniques. 3 U -U -aI., aI., aI., Oxygen Isotopes Traditionally thought to be the result of crustal assimilation (Taylor, 1968), variation both in elemental concentrations and oxygen isotope compositions in igneous systems has recently been proposed to reflect primarily variation in the source of magma (Taylor & Sheppard, 1986; Annen et al., 2006b). If so, subsequent ascent and emplacement of the magma will not significantly alter the chemical or isotopic makeup • • • • rv» 18 of the magma. Hence, if magma increments are derived from sources with different 6' °0 18 values, the increments will have distinct 6 O values. Differences in 6 O values are likely to be preserved during emplacement and crystallization unless modified by local assimilation or subsequent subsolidus (hydrothermal) processes. High temperature processes such as partial melting in source regions and subsequent crystallization of magma will be associated with correspondingly small oxygen isotope fractionations. Hence the variations in the o 1 8 0 value of individual magma increments may be small. Although oxygen isotopes have been used for some time in investigating igneous processes (e.g., Taylor, 1968, 1986), more recent improvements in analytical precision allow documentation of subtle differences in isotope values. Laser fluorination techniques can routinely produce measurements of 61 8 0 with error of +0.10%o (Valley, 2001). This analytical technique was used to measure 51 8 0 values of both whole rock and mineral separates from samples of the Alta stock to attempt to document the presence or absence of discrete magma increments within the Alta stock. It is possible, however, that some degree of oxygen isotope exchange has occurred below the solidus. Previous work (e.g., Kemp, 1985; Bowman et al., 1985; Chen 4 Isotopes aI., of the magma. Hence, if magma increments are derived from sources with different 0180 values, the increments will have distinct 0180 values. Differences in 018 0 values are likely to be preserved during emplacement and crystallization unless modified by local assimilation or subsequent subsolidus (hydrothermal) processes. High temperature processes such as partial melting in source regions and subsequent crystallization of magma will be associated with correspondingly small oxygen isotope fractionations. Hence the variations in the ()ISO value of individual magma increments may be small. Although oxygen isotopes have been used for some time in investigating igneous processes (e.g., Taylor, 1968, 1986), more recent improvements in analytical precision allow documentation of subtle differences in isotope values. Laser fluorination techniques can routinely produce measurements of () 180 with error of ±0.1 0%0 (Valley, 2001). This analytical technique was used to measure ()ISO values of both whole rock and mineral separates from samples of the Alta stock to attempt to document the presence or absence of discrete magma increments within the Alta stock. It is possible, however, that some degree of oxygen isotope exchange has occurred below the solidus. Previous work (e.g., Kemp, 1985; Bowman et aI., 1985; Chen 5 et al., 2007) has shown that minerals can equilibrate to subsolidus temperatures in cooling plutonic rocks, including the Alta stock. Even if this is the case, valuable information can still be garnered about early emplacement related processes. In addition, the suite of mineral separates analyzed can constrain the temperature conditions and degree of oxygen isotope exchange during the early cooling history of the pluton. TitaniQ/Cathodoluminescence TitaniQ (Ti-in-quartz) thermometry (Wark and Watson, 2006) and cathodoluminescence (CL) imaging are techniques that show promise if used in concert. CL imaging reveals textures not visible in plane polarized light that are being increasingly used to interpret processes of grain growth and recrystallization. Previous studies have used CL in detrital quartz to establish both provenance and diagenetic history (Gotze et al., 2001) and in metamorphic quartz to establish metamorphic grade in quartzites (Sprunt et a l , 1978), but it is only recently that CL imaging has been applied to igneous systems (e.g., D'Lemos et al., 1997; Muller et al., 2002). For the ends of this study, CL images serve a twofold purpose: 1) to observe textures in quartz not visible in plane light; and 2) to guide locations of Ti analyses within quartz grains so that the resulting TitaniQ temperatures can be interpreted within a meaningful textural context. Wark and Watson (2006) demonstrated experimentally that the Ti content of quartz increases regularly with increasing temperature at a fixed activity of TiC>2. In calibrating the TitaniQ thermometer, Wark and Watson (2006) used CL images of their quartz grains as a crosscheck for titanium content. Though CL is incompletely understood (Gotze et al., 2000) a good correlation exists between CL intensity and Ti content (Wark and Spear, at, sub solidus TitaniQICathodoluminescence aI., aI., aI., aI., Ti02• aI., 6 2005). Since Ti is a luminescent element, higher Ti concentration produces brighter CL domains. Both theoretical and experimental evidence shows that Ti content correlates positively with quartz crystallization temperature (Wark and Watson, 2006). The TitaniQ geothermometer has been applied to two well-known igneous systems, the Bishop Tuff (Wark et al., 2007) and the Vinalhaven Granite (Wiebe et al., 2007) to test the validity of the method. In the Bishop Tuff, CL imaging reveals regular core-to-rim (CR) zoning in quartz phenocrysts, and a bright CL rim. Such a pattern suggests a period of crystal growth in a magma during declining temperatures, then a short period of higher temperature growth prior to eruption. Both TitaniQ temperatures and CL images corroborate this history of grain growth and late-stage thermal rejuvenation just prior to eruption. The Vinalhaven Granite is known to have a late stage input of mafic magma based on independent evidence, and this event is recorded in bright CL rims in quartz grains from the surrounding material. While the TitaniQ geothermometer shows promise, it has not to date been applied rigorously to a typical granodiorite pluton. Different emplacement processes should leave behind different geochemical signals. Distinct temperature patterns would be predicted in an intrusion produced by single-stage vs. incremental emplacement (Annen et al., 2006a). It is expected that the crystallization histories of individual quartz grains would differ depending on the style (single-stage vs. incremental) of pluton assembly. Quartz grains should show a simple pattern of decrease in Ti concentration from core to rim (e.g., declining temperature) if an intrusion is emplaced in a single stage. If, however, incremental emplacement is responsible for assembly of the igneous body, Ti aI., aI., aI., 7 concentration profiles of individual quartz crystals should show one or more stages of thermal rejuvenation from core to rim, due to reheating from new batches of magma. Geologic Setting of the Alta Stock and Emplacement Models The Alta Stock in the Wasatch Mountains of Utah provides an excellent location in which to test the integrated application of oxygen isotopes, CL imaging and TitaniQ geothermometry techniques to investigate the process of pluton emplacement. It is one of the most well studied plutons in North America (Wilson, 1961; John, 1989; Hanson, 1995; Vogel et a l , 2001; Didericksen and Bartley 2003). The Alta contact aureole is also exceptionally large and well studied (Moore and Kerrick, 1976; Kemp and Bowman, 1984; Bowman et al., 1994; Cook and Bowman, 1994; Cook et al., 1997; Chadwell and Bowman, 2005; McLin et al., 2008). As such, a wide variety of data are available with which to integrate oxygen isotope and TitaniQ geothermometry results from the Alta stock. The Alta Stock (Fig. 1.1) is a mid-Tertiary granodioritic pluton that is exposed in the central Wasatch Mountains east of Salt Lake City, UT. It is one in a series of 11 stocks, which lie along an east-trending belt that is coincident with the Uinta arch (Hanson, 1995). The Uinta arch is a suture between the Archean Wyoming Craton and the Proterozoic Yavapai province; it is possible that the suture between these two crustal blocks provides some degree of spatial control for the Wasatch Intrusive Suite. The stocks are between 36 and 30 Ma in age, and follow a high-potassium and calc-alkaline chemical trend (Crittenden et al., 1973; Vogel et al., 2001). The stocks are more deeply exposed in the west and exposure becomes shallower to the east, with magmatism Models aI., aI., aI., aI., aI., H9 B6 Tertiary Intrusions Mafic Dikes Central Phase* n f A l t a S t o c k Border Phase J ] Clayton Pk. Stock Sedimentary Units ^]Quat. Alluvium ] Miss. Carbonates ] Cambrian Clastics N D3 89-1-6 JH6 C6 C5 89-1-11 B4 B3 D2 JB2 88-I-7 H4 H3_ H2 89-I-9 1 km ft •HI I Figure 1.1. Location (inset) and simplified geologic map of Alta Stock area showing sample site locations. Border phase (B) and central phase (C) samples are granodiorite, as are samples labeled 88- or 89-. Other samples sites designated by H (hammer site) and D ( drill site) have one or more phases sampled per site (e.g., H2E2). See text for unit descriptions. Map after Baker et al. (1966) 1 PhaSe} D Alta Stock D Sedimentary Units D D D UTAH 00 eventually being manifested by the Keetley Volcanic field east of Park City, UT (John, 1989). The stocks are small in outcrop area (5-100 km2), and can be divided into eastern and western groups based on chemical makeup and texture (Hanson, 1995). The western stocks (Little Cottonwood, Alta, and Clayton Peak) are more felsic and phaneritic than the eastern stocks (Mayflower, Valeo, Pine Creek, Ontario, Flagstaff Mountain, and Glencoe) which are more mafic and generally porphyritic in texture. The Clayton Peak stock is likely the oldest body is the group, while all other stocks and the Keetley Volcanics become younger from east to west. The Alta Stock is approximately 10 in area, and intrudes Precambrian and Paleozoic sedimentary units (Baker et al., 1966; Crittenden, 1965). Wall rocks include Precambrian and Cambrian elastics (quartzites) and Cambrian-Mississippian carbonates (marbles and dolostones). Wilson (1961) and Cook and Bowman (1994) estimated the lithostatic pressure at which the Alta stock was emplaced to be 1.0-2.0 kb, or a depth of -3.5-7 km. The dominant minerals in the stock are plagioclase feldspar and quartz. Smaller amounts of potassium feldspar, biotite, and hornblende are present. Accessory minerals are apatite, zircon, titanite, and magnetite. On the basis of thin section study, Wilson (1961) mapped two phases in the stock, a central porphyritic phase and a nonporphyritic border phase. The porphyritic nature is well developed in some thin sections, but a gradient certainly exists between porphyritic and nonporphyritic phases. Mafic enclaves are found throughout, although these are more common on the outer border of the stock, especially in the western and deeper end of the stock. Enclaves contain the same minerals with the same chemical compositions as the main granodiorite; enclaves simply have a higher proportion of mafic minerals (Hanson, 1995). The contact km2 ), Mayf1ower, Val eo, 9 km2 aI., clastics ~3.5-7 10 between the central and border phases is clear on the southern mapped boundary, but is much less clear on the north and in the west. The stock has been tilted to the east by Cenozoic faulting (John, 1989). As a result, the southwestern portion of the stock is exposed at a deeper level than in the east, which is consistent with the overall trend of the other Wasatch stocks. The stock is cut by the Silver Fork fault in its westernmost part (Crittenden, 1965). Two end-member emplacement models have been proposed for the Alta Stock. Bartley et al. (2008) suggest that the Alta stock was emplaced incrementally as a series of dikes. Field evidence for this hypothesis includes spatial association of the stock with parallel dike swarms, tabular shape of the stock (Fig 1.1), lithologic sheeting within the stock subparallel to the intrusion walls, and synkinematic emplacement indicators (e.g., fractured feldspar grains, biotite kinks, offset dikes). In addition, the Alta stock is likely to have been emplaced in an extensional setting (Vogel et al., 1997, 2001). If tectonic dilation is the primary space-making mechanism, then the -2.5 km-wide Alta stock was emplaced on a 106yrs time-scale given reasonable tectonic rates o f - 1 mm/yr (Bartley et al., 2008). In contrast, Cook et al. (1997) used a geologically instantaneous («5,000 yrs.) emplacement to model heat and fluid flow responsible for the contact aureole into the wall rock surrounding the Alta stock. Their work shows that a short-term influx of advective heat flow is required to produce the exceptionally large lateral extent of contact metamorphism, but limited extent of 1 8 0 / 1 6 0 depletion, seen surrounding the stock. If the Alta stock was emplaced incrementally on the time scale of > 106 years, the amount of heat available for the contact aureole is too low for too long to account for the ai. ai., 1997,2001). ~2.5 106 yrs of ~ 1 ai.,2008). ai. «<5,000 180/160 the If 1 06 11 temperature and extent of metamorphism. The time discrepancy between the two models is orders of magnitude. There are at least three possible solutions to this conundrum: the time scale suggested by the thermal modeling for advective heat flow is as much as 1 OOOx shorter than the timescale of stock emplacement, the stock was emplaced rapidly (<5,000 yrs), or more heat was added to the system from another, unseen body. The Alta stock could be much larger than what is exposed and could provide the requisite heat (Bartley et al., 2007b). However, the geometry of the exposed igneous contact and spatial pattern of isograds in the aureole do not indicate significant lateral expansion of the stock with depth. Both oxygen isotope and Ti temperature data from the Alta stock could give valuable insight into the relationship of time, temperature, and space as it relates to the Alta stock and its contact aureole. Some preliminary oxygen isotope analyses have been made (Kemp, 1985; Taylor, 1968) in the Alta stock. Reported values have small, but significant variations in both whole rock (>l%o) and quartz (~1.5%o). It is unknown, though, how or if these values correlate with lithology and/or location within the stock. 1000x shorter than the timescale of stock emplacement, the stock was emplaced rapidly «5,000 aI., 1%0) (~1.5%0).1t CHAPTER 2 METHODS Sample Collection, Preparation, and Analytical Procedure Field work was undertaken in the summer of 2008. Samples were collected on a pluton-wide SE-NW transect across the stock every -100-200 m (Fig 1.1). All samples from this transect are the dominant granodiorite phase. The purpose of this transect was to try to detect any large-scale variation across the border and central phases, as well as within each phase. Samples were also taken at the contact of the border and central phase (for example, samples H6B1, H6B2, H6C1, H6C2). A series of samples was also taken around the outer margin of the stock. The most lithologic variation occurs near the outer margin of the stock and samples were collected to represent every rock type (i.e., granodiorite, mafic enclave, aplite, etc.) present at each border site location (Fig. 2.1). A detailed drill core transect across a mafic band (Fig. 2.2) was made at sample site D3 (Fig. 1.1). Banded units are present throughout the southern margin of the stock, and were also observed in the eastern and northern margins (Fig. 2.3). All samples were collected from fresh outcrops free of visible hydrothermal alteration or weathering. Procedure ~ 100- 200 H6B1 , H6Cl , Figure 2.1. Outcrop photos of sample sites H2 & H3 showing lithologic variation common around the outer margin of the stock. At sites such as this, samples were taken of most or all rock types (granodiorite, mafic, aplite) present. 14 10 cm Figure 2.2. Photos of sample site D3. Top: Outcrop photo of site D3 showing mafic band within granodiorite and drill core holes . White arrows/numbers mark intervals used in the preparation of isotope samples across this band. Below: Labeled drill cores showing dip of boundary between granodiorite and the mafic band. Figure 2.2. Photos of sample site D3. Top: Outcrop photo of site D3 showing mafic band within granodiorite and drill core holes. White arrows/numbers mark intervals used in the preparation of isotope samples across this band. Below: Labeled drill cores showing dip of boundary between granodiorite and the mafic band. Figure 2.3. Example of a banded feature at site H7 similar to, though narrower than, the mafic band at site D3. These banded features, though small in area, are widespread around the outer margin of the stock. microscope to determine mineralogy, textures, and any unusual features of note. Quartz grains in each section were selected for cathodoluminscence (CL) imaging and electron microprobe (EMP) analysis. Grains were selected to include the full range of quartz morphologies within a sample (i.e., small vs. big, euhedral vs. subhedral, phenocryst vs. matrix, etc.). EMP analyses were done at the University of Utah and at the University of New Mexico (UNM) on a Cameca SX-50 and a JEOL 8200 Electron Probe Microanalyzer, respectively. The Utah probe was run at 150 nA of current with a spot size of 10 um and a count time of 360 seconds. The UNM probe was run at 50 nA of current with a spot size of 2 \xm and a count time of 360 seconds. Ti content was measured on each probe using a PET crystal. One crystal was used at UNM, while two were used simultaneously at Utah. Images of thin sections with probe sites are compiled in the CD insert (Appendix A). JEOL JSM-5800LV cm 16 At least one thin section was prepared from each hand sample. The thin section location was chosen to be representative of the hand sample as a whole. If a hand specimen was heterogeneous or contained multiple petrologic domains (e.g., granodiorite and mafic enclave), additional sections were made to include these domains and petrologic boundaries. Sections were then optically examined using a petrographic lEOL !-tm current with a spot size of 2 !-tm and a count time of 360 seconds. Ti content was measured on each probe using a PET crystal. One crystal was used at UNM, while two were used simultaneously at Utah. Images of thin sections with probe sites are compiled in the CD insert (Appendix A). Cathodoluminscence (CL) images of quartz grains were made at UNM on a lEOL lSM-5800L V scanning electron microscope (SEM) to document CL textures in quartz grains and to guide EMP analyses. Mineral separates and whole-rock powders were prepared for oxygen isotope analysis. Whole-rock samples were ground to a uniform powder of less than 200 mesh using a shatterbox. Preparation of mineral separates began with crushing of about 5-6 cm3 through a Frantz magnetic separator to split the sample into a quartz-feldspar separate and a mafic separate. The quartz-feldspar separate was split into two fractions. One fraction was treated with hydrofluoric acid (HF) for 9-15 minutes to remove feldspar and impure quartz grains. Resulting quartz separates were then inspected under a binocular microscope and any remaining impure grains were picked out by hand. Feldspar (plagioclase+K-feldspar) was separated from quartz in the second fraction by density using the heavy liquid sodium polytungstate. Magnetite, biotite, and hornblende separates were hand-picked for some samples as well. O2. CO2 a Finnegan MAT 251 magnetic sector mass spectrometer. Samples were run against the UWG-2 standard (Valley et al., 1995), giving a precision of between 0.01-0.06 %o (1 SD). 17 of fresh sample by hand in a steel mortar and pestle. The crushed sample was sieved into a series size fractions. The 80-100, 100-150, and 150-200 mesh sizes were cleaned in an ultrasonic bath with distilled water, ethanol, and acetone to remove dust. A hand-held magnet removed magnetic grains and any metal fragments. These clean samples were run poly tungstate. Oxygen isotope analyses of mineral separates and whole rock powders were done at the stable isotope laboratory at the University of Wisconsin-Madison (UW). Samples were ablated with a laser in the presence of bromine pentafluoride to liberate 02. The resulting O2 gas was purified through a series of liquid nitrogen traps, converted to C02 with a heated (800°C) carbon rod, and oxygen isotope compositions were measured with aI., 0 ANALYTICAL Stable Oxygen Isotopes 6 1 8 0 predicted by theoretical considerations (Kieffer, 1982) and confirmed by experiment (summarized in Chacko et al., 2001); the pattern is qualitatively consistent with exchange 3.5) 500°C 650°C exchange equilibrium with the three other minerals in a number of samples (Figs. 3.3, 3.4, 3.5). Feldspar is both enriched and depleted in 1 8 0 / 1 6 0 relative to equilibrium values with respect to quartz, biotite and magnetite (Fig. 3.5). The deviations from expected equilibrium values are small (<0.80 %o) but exceed the estimated analytical error of +0.1 %o. Most of the nonequilibrium feldspars are from samples in the border phase. CHAPTER 3 ANAL YTICAL RESULTS Isotopes Oxygen isotope data for the mineral and whole rock samples are compiled in Table 3.1 and illustrated in Figure 3.1. Minerals have <")180 values in the relative order aI., equilibrium between the minerals. In order to assess exchange equilibrium more closely, fractionation factors between mineral pairs were plotted (Figs. 3.2-3.S) and compared to equilibrium fractionation factors (Appendix B). Quartz (Qtz), magnetite (Mag), and biotite (Bt) (Figs. 3.2-3.4) approximate exchange equilibrium in the subsolidus temperature range of Sooac to 6S0aC (Figs. 3.2-3.4). Feldspar (Fsp), however, is out of 3.S). 180 / 160 relative to equilibrium 3.S). «0.80 0) ±0. 3 . 1 . (CP). Note that all aplite 6-notation. S a m p l e # WR Qtz Fsp Bt Mag Hbl AQtz-Fsp AQtz-Bt AQtz-Mag AQtz-Hbl AFsp-Bt Granodiorite 1-1-89-1-Bl 4.09 4.06 3.56 H2Bla GD Avg. * 3.45 Table 3.1. Oxygen isotope data listed by rock type: Border Phase (BP) and Central Phase CP). and mafic samples are from the BP. WR (whole rock), Qtz (quartz), Fsp (Feldspar), Biotite (Bt), Mag (Magnetite), and Hbl (hornblende) are in C)-notation. Sample # WR Hb.1 ~Qtz-Fsp ~Qtz-Bt ~Qtz- Mag ~Qtz- Hbl ~Fsp-Bt Border Phase 88-I-7 9.51 88-I-9 7.53 10.02 8.25 5.37 1.77 4.65 2.88 89+9 8.97 B1 7.47 9.21 7.98 4.03 1.23 5.18 3.95 B3 9.13 B4 7.52 9.17 7.93 3.84 1.54 5.82 1.24 5.33 7.63 3.35 4.09 B5 7.38 8.86 7.98 5.01 0.88 3.85 2.97 B6 9.18 B7 7.38 9.51 8.39 4.57 1.12 4.94 3.82 D2B2 7.86 9.39 8.41 4.66 0.98 4.73 3.75 D2B3a 7.53 9.78 8.32 4.91 1.77 4.87 3.41 D2B3b 7.52 9.41 8.46 5.09 1.77 4.32 3.37 D2L1 9.19 D3I1 7.82 9.31 8.23 4.17 1.71 1.08 5.14 7.60 4.06 D3I1 (A) 7.51 D3I10 7.16 9.49 8.00 4.47 1.49 5.02 3.53 D3I2 7.72 9.45 8.60 4.76 0.85 4.69 3.84 D3I3 7.75 9.16 8.21 4.65 1.70 0.95 4.51 7.46 3.56 D3I9 7.93 8.97 8.66 4.68 0.31 4.29 3.98 H2B1a 9.38 H3B1 7.69 9.74 8.31 5.03 1.43 4.71 3.28 H4B1 9.38 H4B2 9.46 H6B1 7.37 8.99 7.98 3.89 1.01 5.10 4.09 H6B2 7.64 8.92 8.10 3.85 0.82 5.07 4.25 H8B1 7.53 9.19 8.24 4.58 0.95 4.61 3.66 GO Avg.* 7.54 9.31 8.21 4.52 1.65 1.12 4.50 7.56 3.45 ..... \0 cont. Sample # WR Qtz Mag AQtz-Fsp AQtz-Bt AQtz-Mag AQtz-Hbl AFsp-Bt Al 8.75 9.34 9.02 4.20 0.32 5.14 4.82 H1A1 9.61 D3I8 8.66 8.96 8.76 4.47 0.20 4.49 4.29 D3I8 (A) 8.57 Aplite Avg. Mafic D2E1 8.66 9.30 4.20 0.26 4.82 4.56 D3I4 9.19 8.53 4.67 1.98 5.70 0.66 4.52 7.21 3.49 3.86 D3I4 (A) 6.73 D3I5 7.38 1.12 5.46 4.34 D3I6 7.31 9.37 9.02 4.60 1.74 6.02 0.35 4.77 4.42 D3I7 7.98 9.20 4.84 4.36 4.01 El 8.29 H2E2 6.51 9.23 1.87 H3E2 9.27 Mafic Avg. 7.14 9.21 8.39 5.86 0.87 4.16 BP Avg. 7.60 9.29 8.32 1.02 4.77 7.51 3.84 CI 9.16 CI Pheno 9.08 C3 9.45 C4 7.34 9.14 1.84 5.95 0.85 5.04 C5 9.34 7.92 4.52 1.42 4.82 C6 9.32 8.96 C7 7.37 8.92 0.94 9.09 9.02 0.47 H6C2 8.99 4.18 9.58 9.07 CP Avg. 7.34 9.18 8.04 1.84 0.94 4.88 7.30 3.94 Table 3.1. cant. Sample# Fsp Bt Hbl ~Qtz-Fsp ~Qtz-Bt ~Qtz-Mag ~Qtz-Hbl ~Fsp-Bt Aplite A1 0318 0318 Avg. 8.89 02E1 9.85 0314 6.90 0314 0315 9.31 8.19 3.85 4 .34 0316 4 .60 7.63 3.35 4.42 0317 8.85 0.35 E1 7.36 Avg. 4.49 1.86 4.78 7.42 4.16 BPAvg. 7.60 4.53 1.73 5.85 3.84 Central Phase Granodiorite C1 C1 8.29 4.10 7.30 3.19 4.19 7.33 4 .82 3.40 C6 Pheno 7.98 4.62 4.30 3.36 C7 Pheno H6C1 8.63 9.49 4.46 5.03 4.56 7.30 7.98 3.80 1.01 5.19 89-1-6 89-1-11 CPAvg. 4.26 5.95 3.94 *The average from sample site D3 was used to calculate the BP GD average as to not bias the calculation due to higher sample density at site D3. tv o 5,80 • • • • • • • • • • • • _ • A A • A A A O O A 0 ° A ° o o Q o oo o 0 o o ° A A A • A A A O O O O 0 00 A • o o o • • • - - • • • Aplite Border Phase Granodiorite Central Phase Granodiorite o Wr Qtz A - Bt Mag • Hbl Rock Type 3.1 Qtz (quartz), Fps (feldspar), Bt (Bi), Mag (magnetite), Hbl (Hornblende). The range in each phase has considerable overlap, but averages of WR and Bt differ, as discussed in text. Values display expected order of 180/160 enrichment (Kieffer, 1982). to 11 • • • • ••• •• iI • •• •• • • •• • •• • • •• .._ II•I ~. • I!!L_ ~ • - '. 0 0 .... ................ .... .... .... .... .... .... .... 0 .... ........ .... 0 ........ .... .... .... 0 0 00 00 0 0 0 00 00 0 0 0 00 0 0 0 9 7 ! • 5 - - - - - - - - - - - - - - --- -- - - - 3 "-" -"- • • • • 1 Aplite Border Phase Granodiorite • ••.•... •.... • • .... .... • 0 00 .... 0 0 0 • • - - - - • • Mafic Material -- • Qtz .... Fsp • Mag Hbl Figure 3,1 Oxygen isotope data for whole rock and mineral separates arranged by rock type. Abbreviations are: WR (Whole rock), ofWR 1801160 .N... .. 16.00 1 14.00 AQtz-Mag Figre 3.2. Plot of Qtz-Bt vs. Qtz.-Mag mineral fractionations. A) Equilibrium fractionation factors of Chacko et a l , 1996 and Clayton et al., 1989 (experimental). B) Equilibrium fractionation factors from Bottinga & Javoy, 1975 (empirical). Measured values plot between the two alternatives, defining an equilibrium exchange temperature of 600-650°C ..... a:l I .N... . 16.00 ~----------------------------------------------------------------------------------, 14.00 +---------------------------------------------------------------------------------__1 12.00 +-------------------------------------------------------------------~~--~------__1 10.00 +---------------------------------------------------------~~--------~----------__1 ~ 8.00 +---------------------------------------------~~----------------~--------------__1 6.00 +---------------------------------~~--------_=~=-------------------------------__1 4.00 +-------------~~~~~--~~----------------------------------------------------__1 2 . 00 +-----~~----~------------~------------~------------~------------,-----------~ 3.00 5.00 7.00 9.00 i1Qtz-Mag 11.00 13.00 15.00 aI., aI., tv tv 2.00 1.50 1.00 -I 1 1 1 1 1 1 \ _ AQtz-Figure 3.3. Plot of Fsp-Bt vs. Qtz-Bt mineral fractionations with equilibrium fractionations from Bottinga & Javoy (1975) and Clayton et al. (1989) for feldspars of compositions An30 and An40 (heavy lines). Some Qtz-Fsp-Bt triplets are not in exchange equilibrium. A number of others approach exchange equilibrium in the T range of 500-600°C. ..... a::l I 0- Vl u.. <l 5 . 00 r---------------------------------------------------------------~--------~~--_, 4.50 +------------------------------------------------=--~~------~~----------------__4 4.00 +------------------------------T~----~~----~~~~----------------------------__4 3.50 3.00 2.50 2.00 +---------------------------------------------------------------------------------~ 1.50 +-----------------------------------------------------------------------------------~ ~----------~----------~----------~----------~----------~----------~--------~ 3.00 3.50 4.00 4.50 ~Qtz-Bt 5.00 5.50 6.00 6.50 e uilibrium. a roach exchan e e uilibrium ran e 10.00 T 9.50 9.00 8.50 8.00 cn JH7.50 O < 7 . 0 0 6.50 6.00 5.50 An30 An40 5.00 0.00 0.20 0.40 0.60 0.80 1.00 AOtz-Fsp 1.20 1.40 1.60 1.80 2.00 Figure 3.4. Plot of Qtz-Mag vs. Qtz-Fsp mineral fractionations with equilibrium fractionation factors of Chiba et al. (1996) and Clayton et al. (1989) for feldspar compositions of An30 to An40 (heavy lines). Qtz-Mag pairs approximate equilibrium near 650°C. Three of the five Qtz-Fsp pairs are not in exchange equilibrium at this temperature. 10 .00 9 .50 9 .00 01 ::r:?u! 7.50 I .N... . a 1 7.00 6 .00 0 .20 0 .60 0 .80 /). 1.00 • 2 .00 Fs airs exchan e e uilibrium tern erature. 0.00 0.20 0.40 0.60 0.80 1.00 1.20 1.40 1.60 1.80 2.00 AQtz-Fsp Figure 3.5. Plot of Qtz-Bt vs. Qtz-Fsp mineral fractionations with equilibrium fractionation factors of Chacko et al. (1996) and Clayton et al. (1989) for feldspars of compositions An30 and An40. Border phase samples are denoted with black circles and central phase samples with open circles. Groups with anomalously high and low AQtz-Fsp values indicated in gray areas. Most of the anomalous values are from the border phase and were not included in average calculations. Remaining mineral triplets achieve or closely approach exchange equilibrium in the T range of 550°-620°C. 8 . 00 ~------------------------------------------------~------------------~~--------------~ 7 . oo +-------------------------------------------~~---------------.~----------------------~ 6.00 +------------------------------------J~==========~~~--------------------------~ _ A1 0316 _ -- 03110 5 . 00 +-----------------------~--------~--~~----~~--~--------------~----------_~~rr-a~~ aN5 -0317 _ C5~H 3B1 88-1- 9- ~4.00+-------------~----0-3-1-9--~--~L-~~~--~--------------------------~~---------~_2_B_3_b ____ ~ 3 . 00 +---------------------~~------~------------------------------------------------------~ 2 . 00 +-----------------~----.,--------------------------------------------------------------~ 1 . 00 +-------~--~~-4~------~--------~------~--------~------~--------~-------r------~ 0040 ~ lAO 2 .00 3.5 . ~Qtz-Fsp subsolidus ( 5 2 8 - 677°C). Qtz- Hbl pairs yield the highest equilibrium temperatures in both the border and 2 0 °C cooler ( 6 5 6 °C vs. 677°C) than the central phase. Qtz-Mag pairs also yield relatively 6 3 9 °6 5 5 °granodiorites is Qtz-Fsp. Fsp exchange temperatures were calculated using only samples determined to approach equilibrium based on Fig 3.5. Equilibrium temperatures average 553°C in the border phase and 5 8 3 °C in the central phase. Uncertainties for mineral exchange pairs differ from pair to pair. The error was estimated by performing the sample calculation as above, but assuming a maximum spread in A-values of 0.2%o (double the analytical uncertainty of +0.1 %o). Qtz-Mag appears to be the most sensitive thermometer, with an error o f + 1 2 ° C . Qtz-Bt has an error 1 5 ° C , ± 2 0 ° C , 6 , 8 0 o o. d 1 8 0 6 1 8 0 of Fsp expected as Fsp contains - 6 0 % of the oxygen in the rocks. Some of the Fsp samples are 26 Isotopic exchange temperatures were calculated from mineral pairs determined to approximate equilibrium behavior (Table 3.2) using the same equations (Appendix B) used for Figs. 3.2-3.5. Average temperatures from mineral pairs vary but are all subsolidus (528- 677°C). Qtz- Hbl pairs yield the highest equilibrium temperatures in both the border and central phase granodiorites, although the border phase yielded temperatures on average 20°C cooler (656°C vs. 67rC) than the central phase. Qtz-Mag pairs also yield relatively high average temperatures of 639°C in the border phase granodiorite and 655°C in the central phase granodiorite. Equilibrium temperatures from Qtz-Bt pairs are equal in the central and border phase granodiorite, averaging 553°C and 552°C, respectively. The mineral pair displaying the greatest difference between the central and border phase granodiorites is Qtz-Fsp. Fsp exchange temperatures were calculated using only samples determined to approach equilibrium based on Fig 3.5. Equilibrium temperatures average 553°C in the border phase and 583°C in the central phase. Uncertainties for mineral exchange pairs differ from pair to pair. The error was estimated by performing the sample calculation as above, but assuming a maximum spread in ~-values of 0.2%0 (double the analytical uncertainty of ±0.1 %0). Qtz-Mag appears to be the most sensitive thermometer, with an error of ±12°C. Qtz-Bt has an error of ± 15°C, Fsp-Bt of + 19°C, Qtz-Hbl of ± 20°C, and Qtz-Fsp of ±85°C. Whole-rock (WR) b180 values from all rock types range from 6.73%0 to 8.75%0. The b180 values of whole rock correlate with the b180 values ofFsp (Fig. 3.6), which is ~60% demonstrably out of exchange equilibrium with quartz and biotite (Fig. 3.5). Hence part Table 3.2. Calculated oxygen isotope equilibrium temperatures. See text for discussion of equations used. All isotope values are in %o notation and all temperatures (T) are in °C. S a m p l e AQtz-Fsp Q t z - F s p T AQtz-Bt Qtz-Bt T AQtz-Mag Qtz-Mag T AQtz-Hbl Qtz-HbIT AFsp-Bt F s p - B tT Bl 5.18 526 540 B4 1.24 468 5.33 516 7.63 635 3.35 4.09 528 B5 0.88 606 3.85 638 2.97 645 1.12 4.94 543 D2B2 0.98 560 4.73 559 3.75 559 D3I1 1.08 520 529 7.60 4.06 531 D3I2 621 4.69 562 3.84 D3I3 0.95 573 4.51 7.46 645 3.56 578 H6B1 1.01 5.10 528 H6B2 0.82 5.07 534 4.25 H8B1 573 4.61 568 GD Avg. 4.83 D3I8 0.20 4.49 578 4.29 512 D3I4 0.66 4.52 578 661 3.49 639 3.86 549 D3I5 1.12 508 330 526 661 3.49 BP Avg. 621 7.30 3.19 520 0.94 578 4.30 595 3.36 1.01 548 5.19 525 4.18 CP Avg. 677 0 0c. Sample ~Qtz-Fsp Qtz-FspT ~Qtz-Bt ~Qtz-Mag ~Qtz-Hbl Hbl T ~Fsp-Bt Fsp-BtT B1 1.23 471 5.18 526 3.95 540 B4 1.24 468 5.33 516 7.63 635 3.35 656 4.09 528 B5 0.88 606 3.85 638 2.97 645 B7 1.12 506 4.94 543 3.82 552 02B2 0.98 560 4.73 559 3.75 559 0311 1.08 520 5.14 529 7.60 637 4.06 531 0312 0.85 621 4.69 562 3.84 550 0313 0.95 573 4.51 577 7.46 645 3.56 578 H6B1 1.01 548 5.10 532 4.09 528 H6B2 0.82 638 5.07 534 4.25 515 H8B1 0.95 573 4.61 569 3.66 568 GO Avg. 1.01 553 4.83 553 7.56 639 3.35 656 3.82 554 0318 0.20 1571 4.49 578 4.29 512 Aplite Avg. 0.20 1571 4.49 578 4.29 512 0314 0.66 742 4.52 578 7.21 661 3.49 639 3.86 549 0315 1.12 506 5.46 507 4.34 508 H2E2 1.87 330 Mafic Avg. 1.22 526 4.99 543 7.21 661 3.49 639 4.10 528 BPAvg. 1.05 547 4.83 553 7.48 644 3.42 647 3.90 547 C4 0.85 621 5.04 536 7.30 655 3.19 677 4.19 520 C7 0.94 578 4.30 595 3.36 599 H6C2 1.01 548 5.19 525 4.18 521 CPAvg. 0.93 582 4.84 552 7.30 655 3.19 677 3.91 546 9.80 - m m • - - • • - • * "a IS) Li. X X X o • • • X g 8.SO - X o oo X X X X X X X X X x x X X X JtK X x x - X X • Quartz 7.30 ^ i i i i i 6.40 6.90 7.40 61 8 0 (WR) 7 9 0 8.40 8.90 Figure 3.6 Qtz or Fsp isotope values plotted against WR isotope values. Oxygen isotope compositions of Qtz and WR do not correlate; oxygen isotope compositions of Fsp and WR correlate positively. This suggests that isotopic variation in Fds is responsible for at least some of the variation seen in WR values. - - - --- - - - - - 9.30 - - - -~ - - - - Ci Vl u.. ~ 0 -)( -- - - )( - )( N - )( ..... 8.80 g )( 0 co lo )( )( )( )( )( )( 8.30 )( )( )( M )( )( )( )( )( )( :. )( )( )( 7.80 Feldspar • Quartz )( 7 .30 +------------------r----------------~----------------~------------------r_----------------~~ 6.40 6 .90 7.40 6'SO (WR) 7.90 8.40 8 .90 ofFsp N 00 of the 6 1 8 0 6 1 8 0 clearly disequilibrium feldspar are not included in calculating average whole rock 6 1 80 values. Without samples with disequilibrium feldspar, WR 6 1 8 0 values have a slightly smaller range from 6.73%o to 8.62%o. Average WR values correlate with rock type: granodiorite values are o, mafic values are o, and aplite is higher at o 6 1 8 0 of WR location within the Alta stock in Figure The average 6 1 8 0 value of WR from the border phase (7.60%o) is higher than the average 6 1 8 0 value of WR from the central phase (7.34%o) by a small amount just beyond analytical uncertainty (±0.1 %o). Similarly, the average 6 1 8 0 value of quartz is perhaps slightly higher (9.25%o vs. 9.02%o) in the border phase than in the central phase. Quartz from samples near the margin of the stock (border phase) is a bit more enriched in 1 8 0 / 1 6 0 compared to the border phase as a whole. In 6 1 8 0 6 l s O values of quartz vary from o to o. While the variation is not great, it is significant. d 1 8 0 o o o o 4.41%o 29 ofthe two per mil range in b1S0 values of whole rock samples may be the result of preferential (with respect to the other major minerals) subsolidus exchange of Fsp with a high-temperature hydrothermal fluid, suggesting open-system behavior. In order to evaluate the variation in b1S0 of the Alta stock at the magmatic stage, samples with b1S0 b180 0. 7.47%0, 7.14%0, 8.66%0 (Table 3.3). The blSO values ofWR and quartz from granodiorite are plotted as a function of 3.7. b1SO ofWR 0) b1SO ofWR 0) 0). bl80 0 0) IS0 /160 to the border phase a whole. contrast, quartz from the three samples of the border phase adjacent to the central phase have lower b1SO values relative to quartz from the nearest central phase samples. The b1S0 8.9%0 10.0%0. Biotite b1S0 values range from 3.8%0 to 4.70%0 and are different between the central and border phase granodiorites. On average, biotites from the central phase are about 0.50%0 lower at 4.17%0 compared to border phase samples at 4.41 %0 (Table 3.3). Table 3.3 Average oxygen isotope values for B P and C P granodiorites, as well as aplite and mafic material. Samples with feldspar out of equilibrium were not included in W R average. RockType WR Fsp Bt Mag Hbl Mafic BP CP mat erial. WR Qtz Hbl Border 7.60 9.25 8.19 4.41 1.65 5.82 Central 7.34 9.02 8.08 4.17 1.84 5.95 Aplite 8.66 9.30 8.89 4.20 7.14 9.21 8.39 4.84 1.86 6.02 w o 10.02 Figure 3.7. Outline of the Alta Stock showing average WR (black) and Qtz (red) 6 1 8 0 data from granodiorite samples. WR 5 1 8 0 values are slightly higher on average in the border phase. Qtz values are slightly enriched on average in the border phase. The highest Qtz 6 1 8 0 values are found in the southern margin of the stock. Border Phase Avg. WR= 7.60 Avg. Qtz= 9.25 9.28 o 7.65 7.53 o 9.19 o 9.58 o 0 8.97 9.14 07.34 9.450 0 9.120 Central Phase Avg. WR= 7.34 Avg. Qtz= 9.02 <') 180 <') 180 emiched <') 180 However, the average 6 1 8 0 values of quartz, feldspar, hornblende, and magnetite are equivalent, within analytical uncertainty, between the central and border phases. Hornblende separates are fairly consistent, averaging 5.90%o. On smaller scales (outcrop and hand sample), WR and quartz d 1 8 0 values also vary. At the drill site D 3 (Fig. 3.8), WR 6 1 8 0 values from mafic material range from o to o and quartz from mafic samples from o to o (Fig. Granodiorite from site D 3 has uniform WR 6 1 8 0 values o to o) but quartz 61 8 0 values vary from o to o. Even though the two rock types have significantly different WR 5 1 8 0 values, as expected, there is no significant difference in o1 8 0 values of quartz from the granodiorite and mafic layer. Hence, on the outcrop scale, the 6 l sO values of quartz do not correlate with rock type. To evaluate the isotopic homogeneity within lithologic bands at site D 3 , material from two different portions of isotope zones 1, 4 , and 8 was analyzed. Zones 4 and 8 have sharp contacts with surrounding material and yield reproducible WR isotope values: o and o for zone and o and o for zone Zone is somewhat more variable o and 7.5 l%o), which may not be surprising as it is not a distinct lithologic zone bounded by sharp contacts as are zones 4 and 8. The 6 O values of WR and mineral separates were also analyzed from two different locations within a uniform-looking hand sample at site D2 (Fig. 3.9). These results show that quartz d 1 8 0 values can be significantly different over this cm-scale 9 . 4\%o to o) even if WR and other mineral 6 O values are equivalent across the hand sample. 32 elSO 5.90%0. elSO D3 elSO 6.90%0 7.38%0 9.19%0 9.52%0 3.8). D3 elSO (7.75%0 7.82%0) elSO values vary from 9.16%0 to 9.45%0. Even though the two rock types have significantly different WR elSO values, as expected, there is no significant difference in elSO values of quartz from the granodiorite and mafic layer. Hence, on the outcrop scale, the elSO values of quartz do not correlate with rock type. To evaluate the isotopic homogeneity within lithologic bands at site 03, material from two different portions of isotope zones 1, 4, and 8 was analyzed. Zones 4 and 8 have sharp contacts with surrounding material and yield reproducible WR isotope values: 6.90%0 and 6.73%0 for zone 4 and 8.66%0 and 8.57%0 for zone 8. Zone 1 is somewhat more variable (7.82%0 and 7.51 %0), elSO ofWR also analyzed from two different locations within a uniform-looking hand sample at site 02 (Fig. These results show that quartz e l80 values can be significantly different over this cm-scale ( 9.41 %0 to 9.78%0) even if WR and other mineral e l80 values are equivalent across the hand sample. Figure 3.8. Sample site D3 showing isotope zones (white arrows) and S 1 8 0 values of quartz (red) and whole rock (yellow). Qtz isotope values are variable in both the mafic interior and the border phase granodiorite. WR values are equivalent in the granodiorite and variable in the mafic interior. Zone 8 is aplite, hence the high S 1 8 0 WR value. 33 0180 0180 Figure 3.9. Photograph of sample D2B3; Yellow lines mark sections of slab crushed for isotope analysis. Oxygen isotope values from the two domains are shown for WR (black), Qtz (red), Fsp (green), and Bt(brown). While WR, Fsp, and Bt values are essen­tially equivalent, Qtz values are significantly different between the two domains of an apparently uniform hand sample. 34 35 Cathodoluminescence (CL) Imaging CL imaging was done on a subset of quartz grains from a representative suite of samples. CL imaging was undertaken because this technique has proven to reveal textures within optically continuous quartz grains that are otherwise invisible using the petrographic microscope. It has also been suggested that CL intensity varies due to Ti concentration in quartz (Wark and Spear, 2005). Therefore, it is useful to image grains prior to Ti analysis to provide textural context in which to interpret Ti results. Quartz grains along the north-south transect exhibit two main CL textures (Fig. 3.10): core-to-rim (CR) zoning and tile-like (TL) patterns. CR zoning, which is consistent with growth zoning, is characteristic of quartz phenocrysts in central phase samples (e.g., CI and C4) and in two border phase samples adjacent to the central phase (e.g., B4 and B5). CR zoning is not completely uniform from sample to sample, but CL is consistently brighter in the middle and darker towards the exterior of the grain. A bright CL band is present on the edge of all phenocryst grains imaged. Interestingly, the two central phase samples near the margin of the central phase (C5, CI) do not show clear CR zoning, but are mostly dark in CL. TL patterns are present in all samples, and cut across any CR zoning present. Small, euhedral grains (sample CI, lower right of Fig. 3.10) also show CR zoning, with the CL bands being correspondingly thinner. Groundmass quartz grains from the central phase are mostly uniform in CL, although intensity ranges from dark to bright. Samples from the outer margin of the stock (border phase) show a wide variety of textures (Figs. 3.11, 3.12). Aplite dike material (sample Al, Fig. 3.11) shows subtle TL Imaging Cl C7) C 1, 3.11,3.12). AI, I ) White bars are 200 \xm. Figure 3.10. CL images of quartz grains from pluton-wide transect. Clear core-to-rim zoning appears in central phase samples and border phase samples adjacent to the central phase. Many phenocrysts have a bright outer rim. Note near ubiquity of thin bright TL patterns (annealed cracks?). Multiple bright CL centers, especially evident in sample C 1) are artifacts of imaging at this magni­fication. !-tm. Figure 3.11. images from collected in text. White bars are 200 ^m 37 FIgure 3.11 . CL llllages border phase granodiorite and aplite samples near the outer margin of the stock. Note the variety of textures present (discussed in text) included mottled, tile like, and stretched styles. Arrows indicate textures discussed !!mafic |im 38 Figure 3.12. CL images from sample H3E2. A) is in the dominant granodiorite, B) is the contact between granodiorite (GD) and a mafic enclave( ME), C) is within the mafic enclave. B) displays possible relic CR (line a-b) and a rejuvenation band (RB). White bars are 200 fA,m 39 patterns as well as thick, dark CL bands in the right-hand third of the grain. These dark bands are not visible in plane light or back-scattered electron images and are not current fractures. These may well represent an older generation of quartz-filled veins or annealed fractures within what is now an optically continuous quartz grain. H4B2a is a banded, mostly granodiorite sample from the very margin of the stock. The two CL images from this sample show intriguing patterns. TL patterns are present. Enigmatic, seemingly smeared out or stretched, irregular banding is seen, mostly around the margins of the grains, particularly in the area marked with arrow A. There is also a very dark CL band cutting across the grain in the lower image (arrow B, Fig. 3.11). This dark band cuts across all other patterns in the grain except for the smeared band, and seems to extend beyond the grain boundary. Complicated, irregular zoning is also present in sample H3E2 (Fig. 3.12b). Some border phase samples show extremely complex CL patterns that are not visible under other imaging techniques (Fig. 3.13) including back-scattered electron, secondary electron, or optical petrography. Though CL textures are complex, some general patterns are evident. Especially clear in the two grains from sample D2B3a is a mottled CL texture. In the grains, remnants of bright-to-intermediate CR zoning appear to be interspersed and overprinted with dark CL domains. CL images from drill site D3 show CR, TL, and irregular zoning within single quartz grains (Fig. 3.14). As in sample H4B2a (Fig. 3.11), a quartz grain from Zone 3 contains an irregular dark CL band that cuts across all other domains (Fig. 3.14). CR zoning is well developed in zone 4. Subtle or possible relic CR zoning appears in the upper image from zone 3 and in the leftmost image from zone 1. The right image from 40 Figure 3.13. Three pairs of CL and back scattered electron (BSE) images from a two border phase granodiorite samples. Single, apparently simple grains show complex textures when viewed in CL. White bars are 200 /-lm. Figure 3.14. CL images of Qtz from 3 zones across the granodiorite-mafic band contact. Locations of zones are shown in inset (white is granodiorite, gray is mafic material, and pink is aplitic). Textures include CR (zone 4, 1?), TL in all, and an irregular dark CL band in zone 3. White arrows indicate possible offset bands discussed in text. White bars are 200 um . 1 ?), !-lm 42 zone 1 in Fig. 3.14 illustrates a quartz grain with vertical alternating bands (relic CR zoning?) of bright/dark CL that appear to be offset by a series of small, subparallel dark CL bands (annealed cracks?). TL patterns appear most intense in Zone 4, but are present in all imaged quartz grains from all zones at site D3. TitaniQ Geothermometry TitaniQ geothermometry has proven to be an accurate recorder of volcanic (Wark et al., 2007) and shallow plutonic (Wiebe et al., 2007) igneous processes. As it has not yet been applied to more deeply intruded plutons, this application to the Alta stock will be a useful test of this technique in more typical granitic intrusions. Because CL imaging shows that texturally simple quartz grains under the petrographic microscope can actually be texturally complex, it is important to place Ti measurements and any TitaniQ temperature calculations within the textural context provided by CL images. Temperatures were calculated using the method described by Wark and Watson (2006). The TitaniQ thermometer was calibrated in the presence of rutile, pure TiC>2 (Wark and Watson, 2006). TiC>2 activity in the presence of rutile is 1. Rutile is not, however, a common Ti-bearing mineral in most granitoid bodies. Typically, titanite (sphene) and/or ilmenite are the dominant Ti-bearing minerals. The activity of TiC>2 will be less than one in most granitoid bodies, and previous work has used an activity of 0.6 for rocks with titanite/ilmenite (Wark et al., 2007). The Alta stock contains titanite, and thus a TiC>2 activity of 0.6 is justifiable. TiC>2 activity is also assumed to be constant over the emplacement history of the stock. Though it is difficult to empirically constrain the activity of TiC>2, measured concentration of TiC>2 is constant in biotite and Geothermometry aI., aI., Ti02 Ti02 l. Ti02 aI., Ti02 Ti02 Ti02, Ti02 43 hornblende throughout the stock (Hanson, 1995), suggesting that assuming constant a(Ti02) is reasonable. With these two assumptions, variations in concentration in quartz can be interpreted as changes in temperature. 3.4. Hanson (1995) estimated the solidus of the Alta stock at ~685°C based on water content of the Alta stock of between 3.25-4.5 wt% during its crystallization. This suggests that any Ti temperatures above 685°C record magmatic conditions. Additional temperature estimates from Hanson (plagioclase-hornblende, zircon and apatite saturation) all yield temperatures above 700°C. In contrast, two-feldspar temperatures yield subsolidus temperatures (600°C), similar to oxygen isotope temperatures from this study. TitaniQ temperatures range from subsolidus (615-630°C) to magmatic temperatures (700-902°C). Magmatic temperatures are far more prevalent than subsolidus temperatures, indicating that the majority of analyses may be recording formation temperatures. Ti concentrations and corresponding temperature calculations generally correlate positively with CL intensity (bright vs. dark) (Fig. 3.15). Inverse correlations are the CI, immediately towards the interior of the grain. Ti Measured Ti concentrations and calculated temperatures are compiled in Table ~685°C study. TitaniQ temperatures range from subsolidus (615-630°C) to magmatic temperatures (700-902°C). Magmatic temperatures are far more prevalent than subsolidus temperatures, indicating that the majority of analyses may be recording formation temperatures. exception, and appear to be most common near the edges of grains and cracks (Fig. 3.16). Ti temperatures measured in grains with well-developed CR CL zoning are consistent with growth zoning. That is, temperatures are higher in the center of phenocrysts and decrease outwards (Fig 3.16). In addition, the bright CL rim present on many phenocryst grains (e.g., Cl, C4, B4; Fig. 3.16) and in nonphenocryst grains from the border phase (Fig. 3.17, H3E2) records higher TitaniQ temperatures than the dark CL bands immediately towards the interior of the grain. Table 3.4. Measured Ti concentration and calculated Ti-in-quartz temperatures. Sample#/Description Site # Grain Spot Ti (ppm) Temp (°C) Comments BORDER PHASE GRANODIORITE 88-1-9 Grain 1, rim-core-rim Sitel 1 1 83 788 1 2 65 758 1 3 67 761 2,3 are small grains 2 1 82 786 3 1 55 737 Average 70 766 BJ-AS08-B4 Site 1 1 1 74 776 Grain 1 is a larger, eul Utah Probe 1 2 73 774 1 3 61 753 Small 2 1 72 773 Small 3 1 53 735 UNM Probe 4 108 823 5 203 918 6 118 836 7 103 817 8 193 910 9 175 894 11 4 501 13 101 814 14 82 786 15 69 764 16 32 677 17 81 785 18 58 743 19 75 775 Site 2 1 1 76 780 Site 2 has a mafic/felsic contact 1 2 71 771 Grain 1 is an elongate (interstitial?) 1 3 69 767 grain. 2 1 63 757 44 (0C) Site 1 a eut Probe UNMProbe (interstitial?) gram. 45 Table 3.4. cont. BJ-AS08-D2B3a Sitel 1 1 68 762 Transect rim-core-rim 1 2 37 691 1 3 20 630 1 4 144 865 1 5 90 798 Site 2 Rim-core 1 2 72 770 1 3 57 741 Site 3 1 1 113 830 One spot on two small grains 2 1 71 768 D2B3a analyses post Site 2 1 1 99 813 C L 2 15 610 3 36 694 4 16 615 5 85 793 6 34 688 7 8 3 15 487 611 9 18 625 Site 1 avg = 72 749 Site 2 avg = 41 677 Site 3 avg = 92 799 Average 55 711 BJ-AS08-D2B3b Sitel 1 1 83 788 1 2 116 834 1 3 145 865 Transect rim-core-rim Site 2 1 1 69 764 1 2 118 836 1 3 62 751 1 5 83 788 Site 1 avg = 115 829 Sample#/Description Site # Grain Spot Ti (ppm) Temp (°C) Comments (0C) Site 1 CL avg. avg.= avg. Site 1 avg. 46 Table 3.4. cont. A couple small grains 2 1 77 779 Average 105 818 BJ-AS08-D3-8 Long, int. grain Sitel 1 1 67 761 1 2 62 752 1 3 76 776 1 4 43 709 l=Big, subhedral grail Site 2 1 1 41 705 2=ragged looking grain 2 1 74 773 3=extra-big inter. 3 1 79 782 Average 63 751 BJ-AS08-D3-10a 2 small grains Sitel 1 1 99 811 Utah Probe 3 1 111 827 1 61 753 UNM Probe la 42 711 lb 43 712 lc 61 752 3 different grains Site 2 1 1 83 788 Utah Probe 2 1 68 762 3 1 71 768 UNM Probe 1 63 757 2 61 752 3 64 759 4 68 765 5 47 723 6 50 729 3 different grains Site 3 1 1 88 795 2 1 92 801 Sample#/Description Site # Grain Spot Ti (ppm) Temp (°C) Comments 1 3 Site 2 1 1 119 837 Sam2le#iDescri2tion S20t 22m) Tem2 (0C) I BJ·AS08·D3·Site 1 This is from isotope zone 1. 1 =BJ·AS08·D3·Site 1 III UNMProbe Ib I UNMProbe I I Table 3.4. cont. Sample#/Description Site # Grain Spot Ti (ppm) Temp (°C) Comments 3 5 516 4 15 604 Site 3 1 52 730 2 7 541 3 182 901 Average 78 739 CENTRAL PHASE GRANODIORITE 89-1-11 Site 1 1 1 74 777 Core Phenocryst Utah Probe 2 92 804 3 83 790 4 74 776 5 61 752 Rim UNM Probe 1 146 867 2 122 840 4 38 695 5 19 626 6 125 844 7 128 847 8 125 844 9 93 803 10 38 695 11 65 757 12 50 726 Phenocryst 2 1 73 774 Rim 2 85 793 3 89 800 Core 4 88 798 5 50 729 Rim Small, subhedral maxtrix grain 3 1 63 756 Interstitial matrix qtz. 4 1 77 781 Larger interstitial qzt. 5 1 50 730 2 51 732 47 (0C) UNMProbe II 48 (°C) 1 2 92 Average 60 724 BJ-AS08-C4 Sitel 1 1 2 1 2 1 741 1 4 UNM Probe 1 2 64 4 64 822 7 1 1 784 1 2 82 1 74 Table 3.4. cont. Sample#/Description Site # Grain Spot Ti (ppm) Temp (0C) Comments 5 16 610 6 37 692 7 10 570 Site 2 126 845 136 856 3 57 741 4 36 689 5 105 819 6 25 652 7 79 781 8 120 838 9 801 10 51 728 Site 1 57 745 Utah Probe 69 767 3 63 756 55 3 82 788 1 78 782 UNMProbe 103 817 755 3 107 822 95 806 5 755 6 107 105 819 8 100 813 Site 2 79 Utah Probe 788 3 776 49 Figure 3.15. CL image with TitaniQ temperatures for a grain from sample D2B3a. Temperatures correlate with CL intensity. The highest temperatures are found in the brightest CL zones. Black/white spots and text are simply for ease of reading and bear no interpretive significance. White bar is 200 um. fA.m. Figure 3.16 Representative CL images from three samples along the pluton-wide transect with microprobe spots and calculated TitaniQ temperatures. Temperatures mainly correlate positively with CL intensity; opposite correlation can occur, and is most prevalent near grain boundaries and cracks (e.g., C4; spots 6, 10). Phenocrysts show well developed CR (higher-lower T) zoning with a higher temperature, bright CL rim. White bars are 200 urn, N.D.=nondetectable. RepresentatIve !-lm, 50 51 Figure 3.17. CL images with microprobe spots and TitaniQ temperatures from sample H3E2. A. Displays low T in the 'core' of the grain and higher T at the 'rim'. B. Tempera­tures are consistent with growth zoning (spots 3-6) and thermal rejuvenation (spot 2). Spot 3 records the highest temperature in the interior of the grain. Spots 3-6 record temperatures correlated with the intensity of CL banding. Spot 2 corroborates the interpretation that this bright CL zone records a thermal rejuvenation event on the outer edge of the grain. The low temperature at spot 1 is likely an edge or hydrothermal effect. White bars are 200 ^m. rim' . !-lm. 52 The average and maximum TitaniQ temperatures in quartz are plotted as a function of sample location within the Alta stock (Fig. 3.18). Average temperatures were the absence of CL textural context. On the stock scale, average temperatures in quartz are similar in the central and border phases, 768°C vs. 758°C, given analytical uncertainty of +(858°C) (822°C). experienced, or preserved somewhat higher temperatures during its crystallization and can vary by 80°C across a hand sample and by as much as 50°C across a single thin section (Fig. 3.20). 3.1S). calculated using analyses that were not near grain edges, cracks, or other features in CL where effects of diffusion or later subsolidus (hydrothermal) alteration and/or recrystallization could modify Ti content from its original value. Not all grains were imaged with CL, however. On samples not imaged with CL, a series of probe spots (3- 12+) were taken from several quartz grains across a thin section. Only the average temperature from these series was used; no grain-scale interpretations were attempted in the absence of CL textural context. On the stock scale, average temperatures in quartz are similar in the central and border phases, 76SoC vs. 75SoC, given analytical uncertainty of ±10°C. However maximum temperatures on average are 30°C higher in the central phase S5S°compared to the border phase S22°The central phase appears to have emplacement history. Ti temperatures are as variable on the outcrop and hand sample scale as on a stock-wide scale. At sample site D3 (Fig 3.19), the average temperature is higher in the mafic layer (zones 4 and 5) by 30°C to nearly 100°C than in the surrounding granodiorite. Yet in a quartz grain from zone 3, no clear CR zoning is present, and temperatures vary over a comparatively narrow range of 711°C to 765°C. Temperatures SO°Figure 3.18. Outline of the Alta stock with granodiorite sample sites showing average TitaniQ temperatures in black and maximum temperatures at each site in red. On average, temperatures in the border and central phases are similar, but average maximum temperatures are higher in the central phase than in the border phase. Border Phase Avg. Temp=758 Avg. Max Temp=822 837 0762 o o 897 0796 o o (356 724 Central Phase Avg. Temp=768 Avg. Max Temp=858 A. \ »/ \ / 7 51 \ / 818 \ V 3 y \ 8 0 \ / 7 5 6 V\ y ' 761 \ 1 / / 705 Figure 3.19. TitaniQ temperature data and a representative CL image from site D3. A. Average TitaniQ temperatures from isotope zones in site D3. Material from zones 6-9 was not analyzed for TitaniQ temperatures, so they are not shown here as indicated by the dashed line. Mafic material from the modal layer at this site is shown in gray and granodiorite is shown in white. B. CL image of quartz grain from zone 3. C. Same CL image in B., superimposed with microprobe probe spots and TitaniQ temperatures that vary between 711-765°C. A. 10/ /751 55 Figure 3.20. Photograph of sample D2B3, showing locations of thin sections (A and B) outlined in red, and areas within (yellow circles) where quartz grains were analyzed for Ti. Temperatures shown are averages of analyses from multiple grains within each circled site. TitaniQ temperatures (yellow) can vary within a single thin section by up to 50 °C and within the hand sample by up to 80°C. CHAPTER 4 DISCUSSION All three techniques utilized in this study, coupled with field and petrographic observations, provide evidence for a rich, complex emplacement and early subsolidus history of the Alta Stock. It appears as though the CL imaging and Ti temperatures better record emplacement and early cooling history, while CL imaging and oxygen isotope data appear to provide strong evidence for subsolidus exchange and replacement processes operating in the stock. However, some oxygen isotope data can be used to interpret magmatic processes. Implications from the Oxygen Isotopic Compositions of the Alta Stock and its Constituent Minerals All major minerals within the Alta stock have experienced oxygen isotope exchange down to subsolidus temperatures ranging from ~550-650°C. Oxygen isotopes do not, therefore, preserve original thermal signatures from the crystallization of the stock. Results of this study are consistent with both previous work in the Alta stock (Kemp, 1985) and with work in other granitoid plutons (Bowman et al., 1985; Chen et al., 2007). Stock Minerals ~550-650°C. ai., ai., 57 The 6 1 8 0 values of quartz from all rock types vary by 1.4%o, well outside analytical uncertainty (±0.1 %o). There is no correlation of quartz 6 1 8 0 values with rock type (Table 3.2). Average quartz 6 1 8 0 values are equivalent between granodiorite from both the central and border phases, mafic enclaves, and aplites within analytical uncertainty. The granodiorite-both central and border phases-is almost uniform in its oxygen isotopic composition, with W R values varying by only 0.3%o, but the 6 1 8 0 values of quartz from granodiorite vary by 1.1 %o. The 1. l%o variation in the 6 1 8 0 values of Qtz can be the result of several factors in addition to being a record of significant variation in 6 1 8 0 value of Qtz at the magmatic stage-and hence evidence for discrete magma increments. It may be that the observed variations in 6 1 8 0 value of quartz from granodiorite reflect variations in the abundance of quartz in a suite of samples that vary in bulk composition but have approximately the same W R 6 O value. This situation is valid where an isotopically homogenous source region experiences different degrees of melting at high (800-900°C) temperatures. Isotopic fractionations between source rock and melt at these high temperatures are small, so different degrees of melting would produce melts of different bulk composition (i.e., with more vs. less quartz when crystallized) but similar 6 1 8 0 values. To test this possibility, a simple mass balance calculation was performed to model changes in 6 1 8 0 values of quartz as a function of quartz abundance in a closed system (Table 4.1). Starting and final abundances of quartz used in the calculation correspond to those in samples with the highest and lowest measured modes (Table 4.2). Modes of quartz and other minerals in these samples were measured with the QEMSCAN. The 6 O values of minerals in each rock are calculated assuming exchange equilibrium a 6180 0, 0). 18type (Table 3.2). Average quartz 6180 values are equivalent between granodiorite from both the central and border phases, mafic enclaves, and aplites within analytical uncertainty. The granodiorite-both central and border phases-is almost uniform in its oxygen isotopic composition, with WR values varying by only 0.3%0, but the 6180 values of quartz from granodiorite vary by 1.1 %0. 1.1 %0 6180 factors 6180 stage-and 6180 WR 6180 (800-900DC) I sotopic 186180 6180 Table 4.1 Calculated changes in 6 1 8 0 values of minerals as a function of varying Qtz abundance in a closed system of constant 6 1 8 0 value. Mol Min=moles of mineral in sample, X 0 = fraction of O in each sample contributed by that mineral. Variation of 0.4%o is predicted. Starting rock contains 14 vol. % Qtz, corresponding to Qtz abundance measured in sample D2B2. Ending rock corrseponds to sample 89-1-11, with the highest measured Qtz abundance of 26 vol. %. D2B2) cm3) 0 X 0 C a l c u l a t e d cm3) 0 X 0 cm3) 0 X 0 cm3) 0 X 0 4 .1 0180 0180 0= 0 0 of26 Initial (02B2) measured Qtz Fsp Hbl Bt WR Qtz Fsp Hbl Bt Measured 9.39 8.41 5.82 4.66 7.86 Vol (em3) 14.00 73.00 9.00 4.00 Mol Min 0.62 0.71 0.03 0.03 Mol 0 in min 2.00 8.00 24.00 12.00 o in rk from min 1.23 5.65 0.79 0.33 XO 0.15 0.69 0.10 0.04 Calculated mineral isotope values Vol (em3) 18.00 70.00 7.00 4.00 Mol Min 0.79 0.68 0.03 0.03 9.31 8.46 4.87 4.87 7.98 Mol 0 in min 2.00 8.00 24.00 12.00 o in rk from min 1.59 5.42 0.62 0.33 XO 0.19 0.66 0.08 0.04 Vol (em3) 22.00 67.00 6.00 4.00 Mol Min 0.97 0.65 0.02 0.03 9.19 8.34 4.76 4.76 7.98 Mol 0 in min 2.00 8.00 24.00 12.00 o in rk from min 1.94 5.19 0.53 0.33 XO 0.24 0.63 0.06 0.04 Vol (em3) 26.00 64.00 5.00 5.00 89-1-11 Mol Min 1.15 0.62 0.02 0.03 9.03 8.18 4.60 4.60 7.98 Mol 0 in min 2.00 8.00 24.00 12.00 o in rk from min 2.29 4.96 0.44 0.42 XO 0.28 0.61 0.05 0.05 VI 00 Table 4.2. QEMSCAN mineral modes for a representative group of samples. BP (border phase) CP (central phase) M i n e r a l B1 C4 A v g . CP A vg Plagioclase 4 1 . 15 4 1 . 95 Quartz 20.87 Kspar 10.86 Biotite 4 . 39 Clays Chlorite Magnetite Titan ite 0.17 Mineral B4 D2B2 H2E2 (4 89-1-11 BP Avg. (P Avg. 38.53 39.59 44.37 42.12 45.39 38.50 41.15 41.95 22.83 13.46 11.93 6.77 18.17 23.56 13.75 20 .87 18.88 19.14 19.67 10 .86 16.67 20.26 17.14 18.47 Amphibole 8.54 11.02 8.14 29.23 4.91 4.90 14.23 4.91 3.83 2.53 3.94 2.07 5.22 4.39 3.09 4.81 5.54 5.88 7.26 5.86 5.22 4.81 6.14 5.02 0.67 5.28 1.95 1.22 2.60 2.02 2.28 2.31 0.48 1.48 0.58 0.39 0.66 0.69 0.73 0.68 Titanite 0.84 1.05 0.72 0.66 0.26 0.70 0.46 Apatite 0.36 0.56 0.79 0.42 0.39 0.54 0.53 0.47 Epidote 0.18 0.21 0.33 0.34 0.10 0.09 0.27 0.10 ~600°C, corresponding to measured equilibrium exchange temperature indicated by isotope fractionation factors between minerals in most of the samples. The mass balance calculations show that variation in the abundance of quartz from 12 to 24% in a closed system of constant oxygen isotope composition produces a 0.4%o decrease in the 6 O value of the quartz. For the six samples for which both measured mineral modes and • 18 c 18 • mineral 6 O values are available, 6 O values of quartz vary by only 0.4%o (Fig. 4.1). Although there is an apparent correlation between the amount of quartz and d 1 8 0 value in this limited suite of samples, note that the variation in 6 O value of Qtz from all samples of the border and central phases is 1. l%o, greater than can be described by this "amount" • ' 1 8 effect. The mass balance calculations indicate that the observed variations in 6 O values of quartz from the granodiorite cannot solely be the result of closed-system variation in the abundance of quartz in these rocks. Additionally, if quartz exchanged oxygen isotopes with other minerals to different temperatures in different samples, some variation in 6 1 8 0 values of Qtz would result. By 18 • plotting 6 O values of quartz against the measured oxygen isotope fractionation between quartz and biotite (Fig. 4.2), it is evident that there is a correlation between 6 1 8 0 of quartz and A(Qtz-Bt). Linear regression of the data in Fig. 4.2 suggests that exchange of oxygen isotopes to different temperatures between quartz and other minerals in a sample may l o account for up to ~0.25%o of the 1.1 %o variation seen in quartz 6 O values. Alternatively, some of the 1.1 %o variation in 6 1 8 0 values of quartz from granodiorite may result from subsolidus exchange with a high-temperature hydrothermal 60 ~600°C, system of constant oxygen isotope composition produces a 0.4%0 decrease in the 018 0 value of the quartz. For the six samples for which both measured mineral modes and mineral 0180 values are available, 0180 values of quartz vary by only 0.4%0 (Fig. 4.1). Although there is an apparent correlation between the amount of quartz and 0180 value in this limited suite of samples, note that the variation in 0180 value of Qtz from all samples of the border and central phases is 1.1 %0, greater than can be described by this "amount" effect. The mass balance calculations indicate that the observed variations in 0180 values of quartz from the granodiorite cannot solely be the result of closed-system variation in the abundance of quartz in these rocks. Additionally, if quartz exchanged oxygen isotopes with other minerals to different temperatures in different samples, some variation in 0180 values of Qtz would result. By plotting 0180 values of quartz against the measured oxygen isotope fractionation between quartz and biotite (Fig. 4.2), it is evident that there is a correlation between 0180 of quartz and ~(Qtz-Bt). Linear regression of the data in Fig. 4.2 suggests that exchange of oxygen isotopes to different temperatures between quartz and other minerals in a sample may account for up to ~0.25%0 of the 1.1 %0 variation seen in quartz 0180 values. Alternatively, some of the 1.1 %0 variation in 0180 values of quartz from granodiorite may result from subsolidus exchange with a high-temperature hydrothermal o Q 8.00 00 7.50 7.00 6.50 6.00 10 10.00 9.50 9.00 8.50 N » BP * * r -- CP May account for 0.4%o .00 12.00 14.00 16.00 18.00 20.00 22.00 24.00 26.00 28.00 30.00 %Qtz Figure 4.1 Plot of measured o 1 8 0 values of Qtz vs. measured amount of Qtz (black diamonds). The solid line represents the variation in S 1 8 0 value of Qtz produced by variation in the amount of Qtz in a rock system of constant 6 1 8 0 value. Though there is an apparent correlation between the calculated and measured values for Qtz, note that the variation in both border phase (BP, red bar) and central phase (CP, blue bar) is greater than is predicted by the calculation. This 'amount' effect may account for up to 0.4%o of the 1.1 %o variation seen in 6 1 8 0 values of Qtz . ..N.... CI 0 -00 va 8.00 7.50 7.00 6.50 10.00 .. • • • • 0 30.00 bl80 ofQtz ofQtz bl80 ofQtz thc ofQtz b180 thc 0 0 variation seen in b 180 values of Qtz . Figure 4.2 Measured S 1 8 0 values of Qtz plotted as a function of the measured oxygen isotope fractionation between Qtz and Bt. Black line is a linear regression of the data. There is a correlation between the two data sets, suggesting that oxygen isotope exchange between Qtz and other minerals to different temperatures in different samples could produce up to 0.25%o of the variation seen in the o 1 8 0 values of Qtz. Shown are the ranges in 6 1 8 0 values of Qtz for the BP (red bar) and CP (blue bar). 10.0 BP 9.5 CP I t .. . .. • , • , 9.0 • • • 8.5 ..N... 8.0 CI 0 7.5 .0...0... May account for 0.25%0 va 7.0 6.5 6.0 5.5 5.0 3.50 3.70 3.90 4.10 4.30 4.50 4.70 4.90 5.10 5.30 5.50 ~Qtz-Bt 0180 ofQtz ofthe 0 0 180 0180 ofQtz fluid. Open-system exchange with fluids is indicated by the reequilibration of quartz, feldspar, biotite, hornblende and magnetite to subsolidus temperatures in the range 550- 650°C. If the minerals of the Alta stock were experiencing closed-system isotopic 3.2-3.5) significant 1985). open-system 6 1 8 0 igneous values for all minerals and whole rocks and no sig nificant , 8 0 / 1 6 depletions 6 1 8 0 isotopic resetting caused by interaction with hydrothermal fluids is difficult to quantify, it is possible that the range in WR 6 1 8 0 values may indicate the amount of alteration. It is therefore possible that ~0.3%o of the o variation in quartz is due to hydrothermal activity. 6 1 8 0 high as 0.9%o. Therefore, only 0.2%o of the measured 1.1 %o variation can be clearly attributed to 6 1 8 0 values of Qtz at the magmatic stage. As this variation is very close to analytical uncertainty, it appears as though that the 6 1 8 0 values of quartz do not clearly identify discrete magma increments in the Alta Stock. This result does not rule out, however, the existence of magma increments. The Alta Stock is a relatively small ( 1 0 km2) intrusion, and it is possible that the source of the stock is isotopically homogenous. If the stock were assembled via the amalgamation of 63 650°C. exchange, mineral fractionations would not plot within expected equilibrium curves (Figs. 3.2-3.5) into the subsolidus region, and would instead record significant disequilibrium effects (Giletti, 1985). The hydrothermal fluid involved in this open­system exchange is not meteoric water. Measured ()ISO values are in the range of normal significant 180 /160 indicative of interaction with low () IS0 meteoric water are detected. While the degree of () IS0 1.1 %0 The combined impact from variations in Qtz abundance, temperature of isotopic exchange, and subsolidus hydrothermal exchange on the () IS0 value of quartz could be as 0.9%0. 0.2%0 1.1 %0 ()180 () 180 10 km2 ) many isotopically equivalent increments, no variation would be detected in the intrusion 64 itself. It is also possible that more refractory minerals such as zircon and titanite would be more resilient to subsolidus exchange and alteration, and could better preserve original isotopic heterogeneities within the magma. WR and other mineral 6 1 8 0 values can give useful insight into subsolidus processes. On average, border phase W R values are about o higher when compared to the central phase. This enrichment could be explained by alteration of both Fsp (illite/clay) and Bt (chlorite), which are also enriched by about o in the border phase. The proportion of feldspar in the border and central phases are essentially equal, and biotite content is subequal. Since biotite content is low, it likely does not affect the 1 o 6 O of W R feldspar. There are two possible explanations for the 1 8 0 / 1 6 0 enrichment in the altered Fsp of the border phase. First, because the border phase is older, feldspar would have more time to equilibrate with any magmatic or high T hydrothermal fluid as the border phase cooled. Isotopic exchange from near-solidus to subsolidus temperatures would drive isotope values of the Fsp upwards. Alternatively, feldspar in the border phase may have equilibrated with a higher 6 1 8 0 fluid infiltrating from wall rocks. Fluid in the wall rock would likely have come from Precambrian/Cambrian elastics. Isotope values of such a fluid would be expected to be in equilibrium with their host rocks and have a value of ~10-o (Cook et al., Interaction with this wall rock fluid would enrich the Alta stock in 1 8 0 / 1 6 0 QEMSCAN mineral mode data (Table 4.2) also support the notion that feldspar in the border phase has undergone more alteration than feldspar in the central phase. The 64 6180 WR 0.25%0 0.20%0 subequai. 6180 value ofWR as much as feldspar. 180/160 6180 clastics. ~ 10-15%0 aI., 1997). 180 ;160 . 65 border phase has a higher proportion of clay minerals, likely the result of greater alteration of feldspars and biotite. Qualitatively, feldspars in thin section from the border phase appear more cloudy and altered. The border phase is also more mafic (biotite hornblende) than the central phase on average (16.82% vs. 10.84%), so alteration of feldspars appears to have been significant enough to outweigh the tendency of mafic minerals to drive WR isotope values down. Igneous Textures Revealed by Cathodoluminescence A wide variety of CL textures are evident within the stock. An interpretation of emplacement and cooling events can be derived from these textures. CR zoning appears to be consistent with growth zoning, as reported in previous work (e.g., Wiebe et al., 2007). Grains with the most well developed CR zoning appear in the central phase and in border-phase samples immediately adjacent to the central phase. All phenocryst grains imaged have a bright CL band around the outer edge of the grain, interpreted as a thermal rejuvenation event. TL patterns are interpreted as annealed cracks. An EDS spectrum taken on a TL pattern from sample 89-1-11 (Fig. 4.3) suggests that the cracks are filled not with pure quartz, but contain K and Al, possibly from fine-grained potassium feldspar. Al is a luminescent element, and could contribute to the bright CL character of the TL patterns. Hence, a mixture of both quartz and K-feldspar could cement fractures in quartz grains. The ubiquity of TL patterns indicates that a late, stock-wide period of cracking and healing occurred. Bright TL patterns may represent late evidence of strain during or just after emplacement, because the TL patterns crosscut all other CL boundaries. TL patterns + Cathodoluminescence aI., AI, I F I J I I s r : f i l «B 7 7 k c n u n t s Curso 3 . 8 2 75 I "• • • i i , | | | , 0 1 2 3 4 5 6 7 8 9 k e V -ull s c a l e S 23 k c o u n t s OursnR 0 4 75 0 1 2 3 4 5 G 7 8 9 k e V Figure 4.3. EDS spectra taken from a quartz grain in sample 89-1-11. A) is a spectrum from a uniform CL band; only Si and 0 are clear peaks, indicating essentially pure quartz. B) is a spectrum from a brighter CL TL crack, notice the appearance of Al and K, possibly the result of fine-grained potassium feldspar. is a spectrum from a brighter CL TL crack, notice the appearance of Al and 67 are consistently brighter than the surrounding material. Due to their small width (1 micron or less) and invisibility in reflected light and back-scattered electron images, microprobe analysis was difficult. In an attempt to analyze the Ti content, beam spot size was reduced to 1 micron, but yielded questionable results (Fig. 3.18c). Analyses using a 1 micron beam size consistently yielded Ti concentrations of 45 ppm for domains of distinct CL brightness/intensity (Table 3.4, samples 89-1-11 spots 2a-c; D3-10a, site 1, spots la-b). Previous microprobe analyses using a larger diameter beam show that these distinct CL domains have different Ti concentrations. The consistency of measured Ti concentrations using the small spot size suggests that small spot size limited the amount of material sampled, therefore putting an upper limit on the amount of Ti measured. Both mottled and irregular CL patterns are interpreted as evidence for grain replacement events. Replacement may either be the result of strain-induced pressure solution and reprecipitation or strain-induced fracturing followed by reprecipitation associated with the infiltration of some later fluid (magmatic or hydrothermal). In both alternatives, earlier quartz is replaced by dark CL (and low Ti) quartz. Replacement involves epitaxial growth of quartz, because all quartz grains that exhibit mottled and irregular CL patterns are optically continuous under the petrographic microscope It is quite clear that in the Alta Stock CL images are complex. Previous work using CL images as textural indicators in igneous quartz have mostly found rather simple, CR zoned quartz grains (D'Lemos et al., 1997; Wark et a l , 2007; Wiebe et al., 2007). These studies investigated a granite from the Jersey complex of the Channel Islands, volcanic rocks (Bishop Tuff), and a bimodal intrusion (Vinalhaven). The Jersey granite has been suggested to have grown with a significant amount of magma mixing. The l ret1ected I brightnesslintensity 1 Oa, 1 a-ofTi aI., aI., aI., growth of quartz crystals from the Bishop Tuff is a short-lived event (1000s of years); the growth of the Vinalhaven is likely more protracted (>1 m.y.) but is associated with a larger input of mafic magma than is the Alta Stock. The CL textures of quartz in the central phase of the Alta Stock appear to be broadly analogous to those of quartz from the Vinalhaven (CR zoning with a bright CL band around the edge of the grain); both intrusions were emplaced in extensional settings. If, however, the Alta Stock grew over a more protracted period of time, it seems likely that more complicated textures would arise, especially in the presence of tectonic strain. The ubiquitous presence of TL and other nonuniform CL textures superimposed on CR zoning in the border phase supports the notion of synkinematic emplacement. The central phase has less obvious strain-related features, though TL patterns indicate the central phase experienced some strain as well. It follows that the border phase should show more evidence of healed damage if it was older and was emplaced syntectonically. Field relations show that the central phase is indeed younger than the border phase, and it contains more euhedral quartz crystals. Although CL images strongly suggest that the stock experienced strain during emplacement, strain indicators may not be obvious throughout the stock for a couple of reasons. Ree and Park (1997) have shown that a crystallizing system experiencing high-strain at relatively low temperatures (near its solidus temperature) during crystallization can erase many of the signs of deformation via grain boundary migration through dislocation creep. Alternatively, rearrangement of material (via fluid flow) in a low-strain environment could erase strain indicators. There is evidence for movement of material, as evidenced by both oxygen isotope equilibrium temperatures and some Ti temperatures. 68 It~ high­strain 69 Recrystallizing material either through strain-induced creep or subsolidus rearrangement of material could erase many strain indicators and internal boundaries within the stock. TitaniQ Temperatures and Early History of the Alta Stock In contrast to the extensive subsolidus exchange experienced by oxygen in the minerals of the Alta stock including quartz, Ti contents of quartz phenocrysts record magmatic temperature conditions. Grains that show the best CR zoning in CL display similar Ti temperature patterns as well (Fig 3.16). In phenocrysts from samples CI, C4, and B4 temperatures are highest in the core of the grain, decrease towards the edge, and then are higher at the edge of the grain. A reasonable interpretation of this pattern is that phenocrysts crystallized out of a cooling magma, and then experienced a thermal rejuvenation event either near the end of their formation or at some time after initial crystallization. Temperatures in the cores of grains range from 800°C up to 900°C, decrease to around 675°C, and are higher in the bright CL edge of the grain, in the range 775-840°C. Although this is the norm, some points within phenocrysts do not adhere to this pattern. For example, point 6 in sample C4 is in the core of the grain, but has a temperature of 650°C (Fig 3.16). This point is near a dark CL feature (crack?), and may be affected by its proximity to such a feature. Most temperatures that do not follow the typical pattern in phenocrysts lie near apparent cracks or near the edges of grains, and may not record primary temperature conditions (e.g., point 8, sample B4; point 10, sample C4). Although true phenocrysts are largely absent from the border phase, relic growth zoning does exist (e.g., samples 88-1-9, Fig. 3.11; D2B3a, Fig. 3.13; H3E2; Fig. 3.17b). Stock sub solidus C1, be affected by its proximity to such a feature. Most temperatures that do not follow the typical pattern in phenocrysts lie near apparent cracks or near the edges of grains, and may not record primary temperature conditions (e.g., point 8, sample B4; point 10, sample C4). 70 Temperatures are highest in the relic cores of grains, ranging from ~750°C to as high as ~900°C, though maximum temperatures around 800°C are more common. Especially obvious in Fig. 3.17b are alternating bands of light and dark CL. It is possible that these bands reflect oscillatory temperature conditions, but the density of electron microprobe analyses in this grain is not sufficient to assess this hypothesis. Given the data, it appears as though temperatures simply decrease from a high of 894°C at spot 3 to a low of 706°C at spot 6. A thermal rejuvenation event is also indicated in this grain by a bright CL band characterized by a high temperature, 814°C (spot 2) that crosscuts other CL domains. Nearly all phenocryst grains analyzed in the central phase and some non-phenocrysts in the border phase also show evidence for a late stage thermal rejuvenation/recrystallization event. Bright CL rims with correspondingly high Ti temperatures are especially well developed in samples C4 and B4 (Fig 3.16). Additional border phase samples (e.g., H3E2, Fig. 3.17b; D3 zone 1, Fig. 3.14) exhibit the same bright, hot rim. The available data indicate that this thermal rejuvenation affected almost all of the central phase along the north-south sampling traverse. The large spatial scale of this reheating requires either pervasive infiltration of hotter magma through/into an existing central phase, or the assembly of the central phase by multiple injections of magma. In the latter case, the higher temperature rims on the quartz phenocrysts record the local thermal impact of these injections. Evidence for localized heating at the outcrop scale may be provided by the Ti geothermometry results at site D3 (Fig. 3.18). Averages of Ti temperatures indicate that the mafic band is hotter than the surrounding granodiorite, and that average temperatures increase across zones 1 to 3 of the granodiorite toward the mafic band. This suggests that ~ ~900°C, non­phenocrysts 03 ofthe 3 .18). the band at D3 is a dike, which intruded the granodiorite and thermally reactivated a cm width of the granodiorite. 71 ~30 ofthe CHAPTER 5 CONCLUSIONS Oxygen Isotopes Oxygen isotopes provide clear evidence that the major minerals of the Alta stock subsolidus 650°C. The WR 6 1 8 0 values of granodiorite are nearly homogeneous across the stock. Quartz 6 1 8 0 values, however, display significant variation (l.l%o), although on average are equivalent between rock types and across the stock. Mass balance calculations show that this ~l%o variation is too large to result from changes in the abundance of quartz within an isotopically homogeneous reservoir or from isotopic exchange with other minerals at different temperatures. It is possible that the range of quartz 6 1 8 0 values record the minimum isotopic heterogeneity of the magma system, even though some amount of subsolidus exchange with hydrothermal fluids has occurred. Such heterogeneities are consistent with, but do not prove the existence of discrete magma increments within the Alta stock. Cathodoluminescence/TitaniQ Temperatures CL imaging and TitaniQ temperatures appear to record events associated with both crystal growth at magmatic temperatures and subsolidus behavior. The presence of Isotopes have experienced oxygen isotope exchange into the sub solidus at temperatures of 550- el80 el80 1.1 %0), ~ 1 %0 el80 record the minimum isotopic heterogeneity of the magma system, even though some amount of subsolidus exchange with hydrothermal fluids has occurred. Such heterogeneities are consistent with, but do not prove the existence of discrete magma increments within the Alta stock. CathodoluminescencelTitaniQ Temperatures complex CL textures necessitates caution in selecting and interpreting spots within a grain for TitaniQ temperature analysis. Spot selection for EMP analysis should always be guided by CL imaging, if possible. CL images record a rich textural diversity otherwise invisible in routine optical petrography. In the Alta stock, several key textures are present: core-to-rim (CR) growth zoning, bright CL rims interpreted as thermal rejuvenation events, and tile-like (TL) patterns. TL patterns are omnipresent, and are believed to be healed fractures. CL and Ti concentration correlate positively except where probe spots are near the edges of grains or near unusual CL features (i.e., cracks, etc). CR zoning is most common in phenocryst grains, though possible relic CR zoning is found in samples in the outer margin of the border phase. TitaniQ geothermometry records temperatures consistent with growth zoning. Crystals began forming at 850- 900°C and continued to grow as temperature decreased to around 685-700°C. Thermally rejuvenated bright CL bands are clearest on phenocryst grains and grains that show relic CR zoning, and record high temperatures of 750-800°C. Thermal rejuvenation bands are most well developed in the central phase along the north-south traverse, suggesting either large scale heat input into the border phase (via melt migration or infiltration) or that rejuvenation bands record local heating events, such as the emplacement of new magma aliquots. TL pattern are ubiquitous and crosscut all other CL domains. Their small size precluded microprobe analysis, but an EDS spectrum suggests they may be filled with quartz and K-feldspar. These patterns give evidence for a late stage (after crystal growth and rejuvenation in some grains) event of grain cracking and healing. 73 All three techniques yield valuable information about the emplacement, crystallization and early cooling history of the Alta stock, and can likely be applied to other plutons. CL and Ti temperatures work well when used in concert. Although oxygen isotope data from this study corroborate the findings of previous work that subsolidus exchange pervades the Alta stock, it is possible that quartz preserves some remnant signatures of the minimum variation in 6 1 8 0 values of magma, and hence, evidence for multiple increments of magma. More refractory minerals (e.g., zircon, sphene) could retain a primary isotope signature reflective of original magma composition. The emplacement of the Alta stock (and likely other granites as well) was a complicated process. It does seem clear that the traditional idea of big tank emplacement does not operate in the Alta stock. Though individual increments cannot be conclusively mapped from this work, the data suggest a protracted history of emplacement. The acquisition of the geochemical and textural characteristics of the Alta stock is the result of the integrated impacts from source melting, crystallization, and early cooling history. order to form a more holistic view of pluton emplacement, one must gather information about the entire emplacement process from source melting through late-stage cooling and high temperature (~550-600°C) fluid alteration. 74 (', 180 multiple increments of magma. More refractory minerals (e.g., zircon, sphene) could retain a primary isotope signature reflective of original magma composition. In (~550-600°C) APPENDIX A THIN SECTION PHOTOS BY SAMPLE SITE 76 The CD insert contains photos of thin sections used for microprobe analysis and CL imaging in this study. Photos are arranged by sample site as classified in the text. Circled/numbered areas on thin sections correspond to site numbers in Table 3.4. Where available, close-up photo mosaics of individual sample sites are provided. APPENDIX B OXYGEN ISOTOPE EQULIBRIUM EXCHANGE FACTORS APPENDIXB 78 Equations used to construct Figs. 3.2-3.5 are as follows, and are based on empirical and experimental measurements from Chacko et al. (2001) and references therein: Exchange Equation Reference(s) AQtz-Mag 6.29(106/T2) Chiba et al., 1989; Clayton et al., 1989 AQtz-An 1.97(106/T2) Clayton et a l , 1989 AQtz-Al -0.18(106/T2) Clayton et al., 1989 AQtz-Fsp (An40) 0.68(106/T2)* AQtz-Fsp (An30) 0.47(106/T2)* AQtz-Am 3.15(106/T2)-0.30 Bottinga and Javoy, 1975 AQtz-Bt = 3.69(106/T2)-0.60 Bottinga and Javoy, 1975 AFsp-Bt (An40) 3.01(106/T2)-0.60** AFsp-Bt (An30) 3.22(106/T2)-0.60** * AQtz-Fsp equations were made using a weighted average of AQtz-An and AQtz-Al **AFsp-Bi equations were made by combining AQtz-Bt and AQtz-Fsp equations. Mineral abbreviations are: Qtz (quartz), Mag (magnetite), An (anorthite), Al (albite), Fsp (feldspar), Am (amphibole), and Bt (biotite). Temperature (T) is in Kelvin in the equations, but is converted to °C in the text. aI. 6Qtz-= 106/T2) 6Qtz-= 106/T2) 6 Qtz-= 106 /T2) 6Qtz-= 106/T2)* 6Qtz-= 106/T2)* 6Qtz-= 3. 15(0.30 6Qtz-= T2)-6Fsp-= 106/6Fsp-= 106/T2)-ReferenceC aI., aI., aI., aI., lavoy, lavoy, 6Qtz-6Qtz-6Qtz-AI 6Fsp-6Qtz-6 Qtz-REFERENCES Akaad, M.K. 1956, The Ardara granitic diapir of county Donegal, Ireland: Quarterly Journal of the Geological Society of London, v.l 12, p.263-290. Annen, C, Scaillet, B., and Sparks, R.S.J., 2006a, Thermal constraints on the emplacement rate of a large intrusive complex: the Manaslu leucogranite, Nepal Himalaya: Journal of Petrology, v. 47, p. 71-95 Annen, C, Blundy, J.D., and Sparks, R.S.J., 2006b, The genesis of intermediate and silicic magmas in deep crustal hot zones: Journal of Petrology, v. 47, p. 505- 539. Baker, A.A, Calkins, F.C.; Crittenden, M.D., Jr, and Bromfield, C.S. 1966, Geologic map of the Brighton Quadrangle, Utah: Geologic Quadrangle Map - U. S. Geological Survey, Report: GQ-0534 Bartley, J.M., Bowman, J.R., Coleman, D.S., and Didericksen, B.D., 2007, The rate problem of emplacement of the Alta stock, Wasatch Mountains: Geological Society of America Abstracts with Programs, v. 39, no. 5, p. 5. Bartley, J.M., Coleman, D.S., and Glazner, A.F., 2008, Incremental emplacement of plutons by magmatic crack-seal, in Petford, N., and Sparks, S. editors, Plutons and Batholiths (Wallace Pitcher Memorial Issue): Transactions of the Royal Society of Edinburgh, v. 97, p. 383-396. Bottinga, Y., and Javoy, M., 1975, Oxygen isotope partitioning among the minerals in igneous and metamorphic rocks: Reviews of Geophysics and Space Physics, v. 13, p. 401. Bowman, J.R., Covert, J.J, Clark, A .H., and Mathieson, G.A., 1985, The CanTung E Zone scheelite skarn orebody, Tungsten, Northwest Territories; oxygen, hydrogen, and carbon isotope studies: Economic Geology and the Bulletin of the Society of Economic Geologists, vol. 80, p. 1872-1895. Bowman, J.R., Willett, S.D., and Cook, S.J. 1994, Oxygen isotopic transport and exchange during fluid flow; one-dimensional models and applications: American Journal of Science, v.294, p. 1-55. l12, e., RS.C., RS.e.; 1.1.1.1.R, 1.R, 1. I-55. 80 Brown, E.H., and McClelland, W.C., 2000, Pluton emplacement by sheeting and vertical ballooning in part of the southeast Coast Plutonic Complex, British Columbia: GSA Bulletin, v. 112, p. 708-719. Chacko, T, Cole, D.R., and Horita, J. 2001, Equilibrium oxygen, hydrogen, and carbon isotope fractionation factors applicable to geologic systems, in Valley, J.W., and Cole, D.R., editors, Stable Isotope Geochemistry: Mineralogical Society of America Reviews in Mineralogy, v. 43, p. 1-81. Chadwell, L.M., and Bowman, J.R., 2005, Crystal size distribution analysis of forsterite in the Alta aureole, Utah: Geological Society of America Abstracts with Programs, v.37, p. 227. Chen, J.F., Zheng, Y.F., Zhao, Z.F., Li, B. Xie, Z., Gong, B., and Qian, H. 2007, Relationships between O isotope equilibrium, mineral alteration and Rb-Sr chronometric validity in granitoids: implications for determination of cooling rate: Contributions in Mineralogy and Petrology, v. 153, p. 251-271. Chiba, H., Chacko, T., Clayton, R.N., and Goldsmith, J.R., 1989, Oxygen isotope fractionations involving diopside, forsterite, magnetite, and calcite: application to geothermometry: Geochimica et Cosmochimica Acta, v. 53, p. 2985-2995. Clayton, R.N., Goldsmith, J.R., Mayeda, T.K., 1989, Oxygen isotope fractionation in quartz, albite, anorthite, and calcite: Geochimica et Cosmochimica Acta, v. 53, p. 725-733. Coleman, D. S., Bartley, J.M., Glazner, A.F., and Law, R.D., 2005, Incremental assembly and emplacement of Mesozoic plutons in the Sierra Nevada and White and Inyo Ranges, California in GSA Field Forum Field Trip Guide, doi: 0.1130/2005.MCBFYT.FFG. Coleman, D.S., Gray, W., and Glazner, A.F., 2004, Rethinking the emplacement and evolution of zoned plutons: geochronologic evidence for incremental assembly of the Tuolumne Intrusive Suite, California: Geology, v. 32, p. 433-436. Cook, S.J., and Bowman, J.R., 1994, Contact metamorphism surrounding the Alta stock: thermal constraints and evidence of advective heat transport from calcite dolomite geothermometry: American Mineralogist, v. 79, p. 513-525. Cook, S.J., Bowman, J.R., and Forster, C.B., 1997, Contact metamorphism surrounding 18 16 the Alta Stock: finite element model simulation of heat-and 0/ O mass-transport during prograde metamorphism: American Journal of Science, v. 297, p. 1-55. Crittenden, M.D, Jr. 1965, Geology of the Dromedary Peak Quadrangle, Utah: Geologic Quadrangle Map - U. S. Geological Survey, Report: GQ-0378. c., l.1.0 1.MCBFYT.+ the Alta Stock: finite element model simulation of heat-and 0/0 mass­transport during prograde metamorphism: American lournal of Science, v. 297, p. 1-55. lr. Crittenden, M.D., Jr; Stuckless, J.S., Kistler, R. W., Stern, T. W., 1973, Radiometric dating of intrusive rocks in the Cottonwood area, Utah: Journal of Research of the U. S. Geological Survey, v.l, p. 173-178. Didericksen, B.D., and Bartley, J.M., 2003, Did kilometer-scale flow of interstitial melt erase evidence of incremental emplacement of the Alta stock? Geological Society of America Abstracts with Programs, v. 35, no. 6, p. 181. D'Lemos, R.S., Kearsley, A.T., Pembroke, J.W., Watt, G.R., and Wright, P. 1997, Complex quartz growth histories in granite revealed by scanning cathodoluminescence techniques: Geological Magazine, v. 134, p. 549-552. Glazner, A.F., 2007, Thermal limitations on incorporation of wall rock into magma: Geology, v. 35, p. 319-322. Glazner, A.F., and Bartley, J.M., 2006, Is stoping a volumetrically significant pluton emplacement process?: GSA Bulletin, v. 118, p. 1185-1195. Glazner, A.F., Bartley, J.M., and Coleman, D.S., 2004, Are plutons assembled over millions of years by amalgamation from small magma chambers?: GSA Today, v. 14, p. 4-11. Gotze, J., Plotze, M., and Habermann, D. 2000, Origin, spectral characteristics and practical applications of the cathodoluminescence (CL) of quartz-a review: Mineralogy and Petrology, v. 71, p. 225-250. Hanson, R.B., and Glazner, A.F., 1995, Thermal requirements for extensional emplacement of granitoids: Geology, v. 23, p. 213-216. Hanson, S.L., 1995, Mineralogy, petrology, geochemistry and crystal size distribution of Tertiary plutons of the central Wasatch Mountains, Utah [Ph.D. thesis]: Salt Lake City, University of Utah, 371 p. Hutton, D.H. W., 1982, A tectonic model for the emplacement of the Main Donegal Granite, NW Ireland: Journal of the Geological Society of London, v. 139, p. 615- 631. John, D.A. 1989, Geologic setting, depths of emplacement, and regional distribution of fluid inclusions in intrusions of the central Wasatch Mountains, Utah: Economic Geology and the Bulletin of the Society of Economic Geologists, v.84, p.386-409. Kemp III, W.M. 1985, A stable isotope and fluid inclusion study of the contact Al(Fe)- Ca-Mg-Si skarns in the Alta Stock aureoloe, Alta, Utah [M.S. thesis]: Salt Lake City, University of Utah, 65 p. 81 lr; 1.I, 1.1.1.1.1., Fe)­Ca- 82 Kemp, W M, III and Bowman, J.R. 1984, A stable isotope and fluid inclusion study of contact Al(Fe)-Ca-Mg-Si skarns in the Alta Stock aureole, Alta, Utah: Geological Society of America Abstracts with Programs, v. 16, p.558, Kieffer, S. W., 1982, Thermodynamics and lattice vibrations of minerals; 5. Applications to phase equlibria, isotopic fractionation, and high-pressure thermodynamic properties: Reviews of Geophysics and Space Physics, v. 20, p. 827-849. Mahan, K.H., Bartley, J.M., Coleman, D.S., Glazner, A.F., and Carl, B.S., 2003, Sheeted intrusion of the synkinematic McDoogle pluton, Sierra Nevada, California: GSA Bulletin, v. 115, p. 1570-1582. McLin, K.S, Bowman, J.R., and Kaszuba, J. 2008, Influence of salinity on t-x(co<2)) stability of talc and the possible role of fluid immiscibility in development of the talc zone, outer Alta aureole, Utah: Geological Society of America Abstracts with Programs, v. 40,p.255 Moore, J N and Kerrick, D M, 1976, Equilibria in siliceous dolomites of the Alta aureole, Utah: American Journal of Science, v.276, p.502-524. Muller, A., Kronz, A., and Breiter, K. 2002, Trace elements and growth patterns in quartz: a fingerprint of the evolution of the subvolcanic Podlesi Granite System (Krusne hory Mts., Czech Republic): Bulletin of the Czech Geological Survey, v. 77, p. 135-145. Pitcher, W.S., and Berger, A.R., 1972, The Geology of Donegal; A Study of Granite Emplacement and Unroofing: New York, Wiley-Interscience, 435 p. Rhee, J.H., and Park, Y. 1997, Static recovery and recrystallization microstructures in sheared octachloropropane: Journal of Structural Geology, v. 19, p. 1521-5126. Sprunt, E.S., Dengler, L.A., and Sloan, D. 1978, Effects of metamorphism on quartz cathodoluminescence: Geology, v. 6, p. 305-308. Stearns, M.A., and Bartley, J.M. 2008, Assembly of the incrementally emplaced McDoogle Pluton, central Sierra Nevada, CA: Abstracts with Programs - Geological Society of America, v. 40, p.92. Taylor Jr., H.P. 1968, The isotope geochemistry of igneous rocks: Contributions to Mineralogy and Petrology, v. 19, p. 1-71. 18 16 Taylor Jr., H.P. 1980, The effects of assimilation of country rocks by magmas on and 8 7Sr/8 6Sr systematics in igneous rocks: Earth and Planetary Science Letters, v. 47, p. 243-254. 1R. AI(16, 1M., AF., x( co<2) ) A, AR., A, A, 1M. Taylor Jr., H.P. 1980, The effects of assimilation of country rocks by magmas on 180 /160 and 87Sr/86Sr systematics in igneous rocks: Earth and Planetary Science Letters, v. 47, p. 243-254. Taylor Jr, H.P. 1986, Igneous Rocks: Isotopic case studies of curcumpacific magmatism, in Valley, J.W., Taylor Jr., H.P., and O'Neil, J.R., editors, Stable isotopes in high temperature geological processes: Mineralogical Society of America Reviews in Mineralogy, v. 16, p. 227-272. Taylor Jr., H.P., and Sheppard. 1986, Igneous Rocks: I. Processes of isotopic fractionation and isotope systematics, in Valley, J.W., Taylor Jr., H.P., and O'Neil, J.R., editors, Stable isotopes in high temperature geological processes: Mineralogical Society of America Reviews in Mineralogy, v. 16, p. 273-318. Valley, J.W., Kitchen, N., Kohn, M.J., Niendorf, C.R., and Spicuzza, M.J. 1995, UWG- 2, A garnet standard for oxygen isotope ratio-Strategies for high precision and accuracy with laser heating: Geochimica et Cosmochimica Acta, v. 59, p. 5223- 5231. Valley, J.W. 2001, Stable isotope thermometry at high temperatures, in Valley, J.W., and Cole, D.R., editors, Stable Isotope Geochemistry: Mineralogical Society of America Reviews in Mineralogy, v. 43, p. 365-413. Vogel, T.A., Cambray, F.W., Feher, L., Constenius, K.N., Copeland, P.C., Flood, T. P., Garzione, C , Gehrels, G.E., Hodkinson, D., Hanson, S.L., Hoist, T. B., John, D. A., Layer, P.W. 1997, Petrochemistry and emplacement history of the Wasatch igneous belt: Guidebook Series (Society of Economic Geologists (U. S.), v. 29, p.35-46. Vogel, T.A., Cambray, F.W., and Constenius, K.N., 2001, Origin and emplacement of igneous rocks in the central Wasatch Mountains, Utah: Rocky Mountain Geology, v. 36, p. 119-162. Wark, D.A. and Spear, F. S.. 2005, Titanium in quartz: cathodoluminescence and thermometry: Geochimica et cosmochimica acta, suppl. 69, p. A592. Wark, D.A., and Watson, E.B., 2006, TitaniQ: a titanium-in-quartz geothermometer: Contributions to Mineralogy and Petrology, v. 152, p. 743-754. Wark, D.A., Hildreth, W., Spear, F.S., Cherniak, D.J., and Watson, E.B., 2007, Pre­emption recharge of the Bishop magma system: Geology, v. 35, p. 235-238. Wiebe, R.A., Wark, D.A., and Hawkins, D.P., 2007, Insights from quartz cathodoluminescence zoning into crystallization of the Vinalhaven granite, coastal Maine: Contributions to Mineralogy and Petrology, v. 154, p. 439-453. Wilson, J.C., 1961, Geology of the Alta Stock, Utah [Ph.D. Thesis]: Pasadena, California Institute of Technology, 236 p. II. curcumpacific lW., lR., lW., 83 lW., l ratio-Strategies lW. O.c., Holst, o. S .. l, Pre­eruption lC.,