And Cretaceous–Paleogene Provenance Evolution in the Utah Foreland: Evidence from U-Pb Ages of Detrital Zircons, Paleocurrent Trends, and Sandstone Petrofacies

Provenance of the Paleogene Colton Formation (Uinta Basin) and Cretaceous–Paleogene provenance evolution in the Utah foreland: Evidence from U-Pb ages of detrital zircons, paleocurrent trends, and sandstone petrofacies

William R. Dickinson1,*, Timothy F. Lawton2, Mark Pecha1, Steven J. Davis3, George E. Gehrels1, and Richard A. Young4 1Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA 2Department of Geological Sciences, New Mexico State University, Las Cruces, New Mexico 88003, USA 3Department of Global Ecology, Carnegie Institution of Washington, Stanford, California 94305, USA 4Department of Geological Sciences, State University of New York, Geneseo, New York 14454, USA

ABSTRACT strata of southern California are compatible an important aspect of provenance analysis that with coeval derivation of arc-derived detritus is not well elucidated by consideration of bulk The fl uviodeltaic Colton Formation (Late in the forearc sands and the Colton backarc detrital zircon populations without close atten- Paleocene–Early Eocene) forms a lobate sand from a common paleodrainage divide tion to constituent subpopulations. Our analysis depositional system that prograded from crossing the Mojave region to connect hinter- is based on U-Pb ages for 4765 detrital zircons the south into the Laramide Uinta Basin of land Nevadaplano and Mexicoplano uplands in 57 sandstone samples, ~8500 fl uvial paleo- northeastern Utah (United States) with a pre- to the north and south. current measurements, and 705 point counts of served sediment volume of ~3000 km3 and a sandstone petrofacies from Cretaceous–Paleo- maximum thickness of ~1000 m. Joint consid- INTRODUCTION gene strata in the Utah–Arizona foreland of eration of detrital zircon ages, paleocurrent the Cordilleran backarc, and U-Pb ages for 905 trends, and sandstone petrofacies permits an The Colton Formation is a succession of detrital zircons in 54 sandstone samples from assessment of Colton provenance relations Upper Paleocene to Lower Eocene fl uviodeltaic coeval strata in the California forearc. in the context of evolving Cretaceous–Paleo- strata exposed in the Roan Cliffs of northeastern In a preliminary study of Colton provenance gene sedimentation in the Utah foreland. Utah (United States) along the southern fl ank of (Davis et al., 2010), it was reported that half the Grains with U-Pb ages younger than 285 Ma the Laramide Uinta Basin. Reaching a thickness detrital zircons in arkosic Colton sandstones derived from the Cordilleran magmatic arc of ~1000 m in Desolation Canyon of the Green have U-Pb ages younger than 285 Ma, implying form ~50% of the detrital zircons in arkosic River, Colton fl uviodeltaic strata interfi nger that headwaters of the Colton dispersal system Colton sand, and were transported ~750 km northward in the basin subsurface with lacus- extended as far south as the Mojave segment of to the Uinta Basin from the Mojave seg- trine strata of the Green River Formation. We the Cordilleran magmatic arc, 750 km or more ment of the arc by the California paleoriver. use U-Pb ages for detrital zircons from Colton south of the Uinta Basin. These U-Pb data dis- Colton sedimentation was the Paleogene and related strata of the Uinta Basin in combina- proved the previous inference (Dickinson et al., culmination of a persistent pattern of Cre- tion with sandstone petrofacies and paleocurrent 1986) that Colton arkosic sand was derived taceous sediment transport northward, sub- trends to reach an appraisal of Colton provenance from Precambrian basement exposed in nearby parallel to the Sevier thrust front, to supple- in the context of evolving Cretaceous–Paleogene Laramide uplifts. The fl uvial mainstem for ment east-directed sediment delivery to the depositional patterns in the Sevier retroarc fore- Colton sediment dispersal from the Cordilleran retroarc foreland from the Sevier thrust belt. deep and Laramide successor basins. arc northward to the Uinta Basin is termed the The ratio of longitudinally to transversely Assessment of Colton provenance provides California paleoriver (Fig. 1) (after Dickin- derived sediment was enhanced in foreland a vehicle to highlight the utility of combining son et al., 2011). We note that Zawiskie et al. strata after Laramide deformation produced information from detrital zircons, petrofacies, (1982) postulated delivery of some signifi cant intraforeland uplifts that screened the fore- and paleocurrents for interpretations of sedi- fraction of fl uviodeltaic Colton sediment to the land belt from Sevier sources. The relative ment provenance and dispersal. We also show Uinta Basin from as far south as central Arizona abundance of arc-derived detrital zircons the value of deconvolving detrital zircon popu- directly east of the Mojave region. that were contributed to strata of the Utah lations into constituent subpopulations that can foreland increased in late Campanian time be treated separately, both on graphical plots MATERIALS AND METHODS and remained high into Eocene time. Detri- and by using Kolmogorov-Smirnoff (K-S) sta- tal zircon populations in Paleogene forearc tistics. The mixing of age subpopulations of Table 1 lists the sandstone samples from Cre- detrital zircons from different provenances in taceous–Paleogene strata of the Utah-Arizona *[email protected]. varying proportions during sediment dispersal is foreland region for which U-Pb ages of detrital

Geosphere; August 2012; v. 8; no. 4; p. 854–880; doi:10.1130/GES00763.1; 21 ﬁ gures; 8 tables; 3 supplemental ﬁ les. Received 9 November 2011 ♦ Revision received 3 April 2012 ♦ Accepted 4 April 2012 ♦ Published online 26 June 2012

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Figure 1. Late Paleocene–Early 110° W Eocene course of the Califor- CA NV NV UT WY nia paleoriver from the Mojave Sevier Ui CO region to the Uinta Basin in thrust relation to Cordilleran geo- front U 40°

paleodrainage e d N

logic features (Nf—Nacimiento i

divide v i

fault, PR—Peninsular Ranges, d P Fig. 3 SN—Sierra Nevada, SNB— F Sur-Nacimento block, WTR— SR western Transverse Ranges) span of Fig. 15 GR Un restored palinspastically after elevated CR 10 Nevadaplano Mvf Dickinson (2011) for Great DD 11 Basin Neogene extension and UT CO e Svf g CC CMB Jacobson et al. (2011) for Neo- a n gene San Andreas slip. Selected GVfb i K TC M N SN ra SJR d HMB detrital zircon samples (DZ) o le (see Table 1): CH (NARCHU), a p 10 (COL10), 11 (COL11), 12 K CR SJ ) CH (COL12), 13 (COL13), CP9, DR (GC (others plotted in Figs. 3 and X 13 CP9 D 15 at larger scales). Laramide EPM 12 M ia M n LC C basins (Maastrichtian–Paleo- o or r R j M lif ve Z a F a ri gene sediment fi ll) after Lawton C o 35° N v M le e a (2008): B—Baca, F—Flagstaff, N o p Salinia f - j S a P—Piceance, SJ—San Juan, a K J v l e B N i TC—Table Cliff, U—Uinta n DR i a r (HMB denotes the structural SNB e g down-bowing of the Henry i o Dvf j M Mountains basin). Laramide o n R i F uplifts (Kelley, 1955): CC— 100 km n Circle Cliffs, D—Defiance, paleodrainage WTR AZ NM K—Kaibab, M—Monument, J paleorivers N—Needles, SR—San Rafael, K d iv Ui—Uinta, Un—Uncompahgre, id e Z—Zuni. Sierra Nevada paleo- submarine rivers and linked submarine canyons and fans PR canyons draining to the delta Laramide P o w a y depocenter (DD) of the Great features Valley forearc basin (GVfb) are CA M e x i c o p l a n o after Dickinson et al. (1979) uplifts BC JV 30° N and Dickinson (2011). Goler- basins Winsett paleodrainage through the El Paso Mountains (EPM) is DZ Colton Formation 110° W after Lechler and Niemi (2011). samples Coastal Paleocene–Eocene Poway (central) and associated subparallel paleorivers near the border of California and Baja California, the downstream Jolla Vieja (JV) submarine fan offshore, and the inland paleodrainage divide are after Kies and Abbott (1982). Blue arrows denote net Eocene paleofl ow (fl uvial) for the Music Mountain Formation (MMF) deposited in paleovalleys and on alluvial aprons south of the modern Grand Canyon (Young, 2001a, 2001b) and for the Mogollon Rim Formation (MRF) toward the Baca Basin (Potochnik, 1989), Cretaceous (Campanian) paleofl ow for the fl uvial Kaiparowits Formation (K) into the Table Cliffs basin (Fig. 15), and Oligocene paleowinds (C) feeding the Chuska erg (Dickinson et al., 2010). Green line denotes approximate boundary between dominantly Cretaceous (K) arc plutons (westward) and dominantly Jurassic (J) and older arc plutons (eastward). Oligocene (post-Laramide) volcanic fi elds: Mvf—Marysvale, Dvf—Mogollon- Datil, Svf—San Juan. Modern Colorado Plateau rivers (dash-dot lines): CR—Colorado (GC indicates the central reach of the Grand Can- yon), GR—Green, LCR—Little Colorado, SJR—San Juan. States: AZ—Arizona, CA—California, CO—Colorado, BC—Baja California, NV—Nevada, NM—New Mexico, UT—Utah, WY—Wyoming.

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TABLE 1. SAMPLES FOR U-Pb AGES OF DETRITAL ZIRCONS FROM UPPER CRETACEOUS AND PALEOGENE STRATA OF THE CORDILLERAN FORELAND (UTAH) AND THE NORTHERN FLANK OF THE MOGOLLON HIGHLANDS (ARIZONA) Number Stratal unit1 Age1 Locality U-Pb data2 3NARCHU Chuska Eocene–Oligocene Chuska Mountains Dickinson et al., 2010 3COL12 Music Mountain Paleogene Peach Springs Wash this paper 3COL13 Music Mountain Paleogene Duff Brown Tank this paper 412JL5 Claron Eocene Table Cliff Plateau Larsen et al., 2010 5COL3 Horse Bench Early Eocene Gate Canyon this paper 5COL4 Green River Early Eocene Gate Canyon this paper 5COL9 “Colton” Paleocene–Eocene San Pitch Mountains this paper 5CO22 Colton Paleocene–Eocene Price River Davis et al., 2009b 5COL1 Colton Paleocene–Eocene Water Canyon Davis et al., 20107 5COL6 Colton Paleocene–Eocene Horse Canyon this paper 5COL7 Colton Paleocene–Eocene Tusher Canyon this paper 5DC3 basal Colton Paleocene–Eocene Tusher Canyon Mathers, 2009 5WADC DeBeque Paleocene–Eocene Piceance Basin Davis et al., 2009b 5COL2 DeBeque Paleocene–Eocene Government Creek this paper 411JL5 Pine Hollow Paleocene–Eocene Powell Point Larsen et al., 2010 49JL5 Pine Hollow Paleocene–Eocene Powell Point Larsen et al., 2010 48JL5 Pine Hollow Paleocene–Eocene Powell Point Larsen et al., 2010 47JL5 Pine Hollow Paleocene–Eocene Powell Point Larsen et al., 2010 5DC4 upper Dark Canyon Late Paleocene Tusher Canyon Mathers, 2009 5DC2 middle Dark Canyon Late Paleocene Tusher Canyon Mathers, 2009 5COL8 lower Dark Canyon Late Paleocene Diamond Creek this paper 5DC1 lower Dark Canyon Late Paleocene Hay Canyon Mathers, 2009 5GRF Flagstaff Late Paleocene Soldier Summit Davis et al., 2009b 4CP1 Canaan Peak Maastrichtian Escalante Canyon Mathers, 2009 46JL5 Canaan Peak Maastrichtian Pine Hollow Larsen et al., 2010 5COL5 North Horn Maastrichtian Willow Creek this paper 5TF1 upper Tuscher upper Campanian Tusher Canyon Mathers, 2009 5MT9 middle Tuscher late Campanian Gray Canyon Lawton and Bradford, 2011 5MT1 lower Tuscher late Campanian Gray Canyon Lawton and Bradford, 2011 5MF4 middle Farrer late Campanian Gray Canyon Lawton and Bradford, 2011 5MF6 lower Farrer late Campanian Gray Canyon Lawton and Bradford, 2011 5MB10 Bluecastle late Campanian Gray Canyon Lawton and Bradford, 2011 44JL upper Kaiparowits late Campanian Blues Overlook Larsen et al., 2010 4KK4 upper Kaiparowits late Campanian Kaiparowits Plateau Lawton and Bradford, 2011 4KK5 upper Kaiparowits late Campanian Kaiparowits Plateau Lawton and Bradford, 2011 4KK6 middle Kaiparowits late Campanian Kaiparowits Plateau Lawton and Bradford, 2011 4KK2 middle Kaiparowits late Campanian Kaiparowits Plateau Lawton and Bradford, 2011 4KK1 lower Kaiparowits middle Campanian Kaiparowits Plateau Lawton and Bradford, 2011 5MN7 Neslen middle Campanian Gray Canyon Lawton and Bradford, 2011 5MN8 Neslen middle Campanian Gray Canyon Lawton and Bradford, 2011 5CP34 Castlegate middle Campanian Willow Creek Dickinson and Gehrels, 2008 3DR Dome Rock succession Campanian Dome Rock Mountains Spencer et al., 2011 6P6GC Grand Castle Campanian Parowan Canyon Johnson et al., 2011 6WF6GC Grand Castle Campanian Webster Flat Johnson et al., 2011 6WP9GC Grand Castle Campanian Paunsagunt Plateau Johnson et al., 2011 42JL5 capping Wahweap early Campanian Henrieville Creek Larsen et al., 2010 4CP40 capping Wahweap middle Campanian Henrieville Creek Dickinson and Gehrels, 2008 4CP39 upper Wahweap early Campanian Henrieville Creek Dickinson and Gehrels, 2008 41JL5* lower Wahweap early Campanian Star Seep Larsen et al., 2010 5CP33 Ferron Turonian Dry Wash Dickinson and Gehrels, 2008 3COL11 Ferron Turonian Caineville Gap this paper 3CP9 Toreva Turonian Black Mesa Dickinson and Gehrels, 2008 3COL10 Dakota Cenomanian Fremont River this paper 5RRR54 Dakota Cenomanian Ruby Ranch Ludvigson et al., 2010 56-8B Buckhorn Aptian Buckhorn Draw Lawton et al., 2010 5CP32 Buckhorn Aptian San Rafael River Dickinson and Gehrels, 2008 5RRR12 Poison Strip Aptian Ruby Ranch Ludvigson et al., 2010 Note: U-Pb ages for samples in italics were determined by thermal ionization mass spectrometry (TIMS) technology at the Australian National University; all other U-Pb ages were determined by common laser ablation–inductively coupled plasma–mass spectrometry technology in the Arizona LaserChron Center at the University of Arizona (asterisk denotes a TIMS sample not used for statistical analyses). “Colton” is a Colton Formation correlative in the San Pitch Mountains. 1See Figure 2 for stratigraphic positions of samples (localities within each stratigraphic unit listed from west to east). 2For U-Pb ages reported in this paper, GPS locations of sample sites, and Excel ﬁles of tabulated U-Pb analytical data are provided in the supplemental ﬁles (see text). 3Geographic locations in Figure 1 (regional map). 4Geographic locations in Figure 15 (Table Cliff basin). 5Geographic locations in Figure 3 (Uinta Basin). 6Exact geographic locations not plotted (abstract reference only). 7Designated as sample COL.

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TABLE 2. DETRITAL MODES OF DETRITAL ZIRCON SAMPLES North Sunnyside delta Colton Music “Colton” Horn (Green River) (main body) Dark Cyn DeBeque Dakota Ferron Mountain Grains COL9 COL5 COL4 COL3 COL1 COL6 COL7 COL8 COL2 COL10 COL11 COL12 COL13 Qm 86 93 55 53 51 53 50 36 67 84 73 47 58 Qp212tr2321266311 Q 88945753535652487390744859

P 73192314141610415117 K 2 tr 17 18 19 19 17 12 14 tr 4 12 16 F 9 3364133333322181 92323

Lvmtrtr213345213126 Lsm3355118112578121712 L 33761411153099152918

Lt5 4 9 6161417421515183019 Note: Preﬁx “COL” refers to Table 1 (and see text). See Supplementary File 1 (see text footnote 1) for sample localities. Modes are based on point counts of 400 QFL (quartz-feldspar-lithics) grains per sample. Monocrystalline grains: Qm—quartz, P—plagioclase, K—K-feldspar. Polycrystalline grains: Qp—polycrystalline quartz (dominantly chert), Lvm—volcanic and metavolcanic lithic fragments, Lsm—sedimentary and metasedimentary lithic fragments. Q—total quartzose grains (Qm + Qp). F—total feldspar grains (P + K). L—total labile lithic fragments (Lvm + Lsm). Lt—total lithic fragments (L + Qp).

zircons are available. For U-Pb ages reported deliberate search for the youngest grains present tion (with P = 1.0 indicating statistical identity). here for the fi rst time, sample localities are (Mathers, 2009). Those samples were used for We conclude that no robust provenance distinc- provided in Supplemental File 11, U-Pb geo- some but not all statistical analyses because the tions can be inferred for two detrital zircon age chrono logical methods in Supplemental File 22, full age spectra of their detrital zircon populations spectra yielding P > 0.05 from K-S analysis. and full U-Pb analytical data in Supplemental are inconsistent with data for other samples from Even where P < 0.05 but age peaks on age dis- File 33, including both concordia diagrams and the same stratigraphic units. tribution curves are the same, 2 contrasting age age-probability plots with age-bin histograms. We use three complementary criteria for spectra may refl ect derivation of detrital zircon Detrital modes for the same samples are pro- comparison of detrital zircon age populations or subpopulations in varying proportions from the vided in Table 2 as a guide to petrofacies. subpopulations. same source rocks within a common provenance. U-Pb ages from most samples (86%) were (1) Age distribution curves in the form of (3) Tabulated age subpopulations (Table 3) determined by laser ablation–inductively coupled age probability plots (Ludwig, 2003) for which of detrital zircons based on best estimates of plasma–mass spectrometry in the Arizona each grain age is cast as a normal distribution grain age ignoring analytical uncertainties. This LaserChron Center at the University of Arizona including its standard deviation of age error, and approach is not statistically rigorous but allows (Gehrels et al., 2008; Dickinson and Gehrels, all the normal distributions for individual grain numerical manipulation to calculate grain age 2009). U-Pb ages for eight samples were deter- ages are then summed into curves normalized to indices indicative of key provenance contrasts. mined by secondary ion mass spectrometry subtend equal areas below the curves. Peaks on Age boundaries between the age subpopulations (SIMS) technology at the Australian National age distribution curves are statistically rigorous are based empirically upon nulls in the patterns University (Larsen et al., 2010). The age spectrum representations of analytical data, but measured of grain ages for the samples of Table 1, but the for one of the SIMS samples is incompatible with age peaks do not necessarily refl ect faithfully peaks for the subpopulations derived from age other data, perhaps because of stratigraphic mis- the actual distribution of grain ages within a zir- distribution curves are better guides to the cen- correlation, and was not used for statistical analy- con concentrate because the randomly selected tral age span of each subpopulation (Table 3). ses of age data (Table 1). U-Pb ages from most grains dated are but a sampling of the total grain The arc-derived subpopulation (I) is delimited samples (91%) were determined for randomly populations. Andersen (2005) showed that age by the maximum age (ca. 285 Ma) of the oldest selected detrital zircon grains, but the dated grains peaks on age distribution curves are dependent igneous assemblages of the Cordilleran mag- from fi ve of the samples were selected preferen- in detail on the random selection of grains dated, matic arc (Dickinson and Gehrels, 2009). We tially as the most limpid and euhedral grains in a as well as on the distribution of grain ages in the refer to older detrital zircons as pre-arc grains. population sampled. 1Supplemental File 1. Word fi le of Detrital Zircon (2) Probability (P) values calculated from STRATIGRAPHIC CONTEXT (DZ) Sample Localities. If you are viewing the PDF Kolmogorov-Smirnoff (K-S) statistics (Press of this paper or reading it offl ine, please visit http:// et al., 1986). To supplement visual inspection of Detrital zircon samples (Table 1) derive from dx.doi.org/10.1130/GES00763.S1 or the full-text article on www.gsapubs.org to view Supplemental age distribution curves, and to shed light on the 20 different stratigraphic units of the Utah fore- File 1. infl uence of random grain selection for the con- land (Fig. 2). Colton samples are among the 2Supplemental File 2. Word file of U-Pb Geo- fi gurations of the curves, we use K-S statistics to richest (~50%) in arc-derived detrital zircons. chronolo gical Methodology. If you are viewing the test the null hypothesis that two detrital zircon PDF of this paper or reading it offl ine, please visit age populations or subpopulations might have Colton Formation http://dx.doi.org/10.1130/GES00763.S2 or the full- text arti cle on www.gsapubs.org to view Supplemen- been selected at random from the same parent tal File 2. population. Where P > 0.05 from K-S analysis, The fl uviodeltaic Colton Formation (Spieker, 3Supplemental File 3. Excel file of U-Pb Geo- with the analytical uncertainties of each grain 1946) of the Uinta Basin underlies the lacus- chronologic Analyses. If you are viewing the PDF age taken into account, one cannot infer with trine Green River Formation (Fig. 3) into which of this paper or reading it offl ine, please visit http:// dx.doi.org/10.1130/GES00763.S3 or the full-text 95% confi dence (0.95 being the inverse of 0.05) Colton strata intertongue laterally and grade article on www.gsapubs.org to view Supplemental that 2 detrital zircon age populations were not upward (Fouch, 1976), and overlies the Flagstaff File 3. selected at random from the same parent popula- Limestone (Spieker, 1946) of lacustrine origin,

Geosphere, August 2012 857

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TABLE 3. AGE SUBPOPULATIONS (U-Pb AGES)1 IN DETRITAL ZIRCON SAMPLES Age subpopulation Age range (Ma)2 Age peaks (Ma)3 Inferred principal ultimate sources4 I. arc-derived grains Cordilleran (USA–Mexico) magmatic arcs Ia. dominantly Cretaceous 60–124 70–100 coastal Cordilleran magmatic arc arc magmatic null5 125–130 no DZ grains of this age range Ib. dominantly Jurassic 131–205 150–180 interior Cordilleran magmatic arc Ic. Permian-Triassic 206–282 240–245 magmatic arcs of northern Mexico

arc-derived/pre-arc division 283–288 ca. 285 Ma no DZ grains of this age range

II. Paleozoic–Neoproterozoic 289–691 Appalachian (–Ouachita) orogen IIa. Paleozoic <540 Ma 410–455 native and accreted magmatic arcs IIb. Neoproterozoic >540 Ma 550–600 accreted arcs and rift plutons

igneous age gap6 692–911 uncertain derivation7

III. Grenville Mesoproterozoic 912–1310 1035–1175 Grenville province (eastern-southern Laurentia)

IV. pre-Grenville Mesoproterozic 1311–1579 1405–1455 anorogenic granitic plutons (southwest Laurentia)

V. late Paleoproterozoic 1581–1855 1695–1740 Yavapai-Mazatzal province (southwest Laurentia)

VI. older Paleoproterozoic 1855–2430 1880–1885 multiple northern Laurentian age provinces

VII. Archean or nearly Archean 2470–30588 2615–2760 Superior province of northeast Laurentia9 Note: DZ—detrital zircons (see Table 1). 1U-Pb ages are based on best estimates of grain ages ignoring analytical uncertainties. 2Ranges adapted empirically from patterns and clusters of U-Pb ages for the DZ samples of Table 1. 3From multiple age distribution curves of this paper. 4Ignoring sand recycling from pre-Cretaceous strata of the Sevier thrust belt or the Mogollon highlands. 5Early Cretaceous interval (minimal duration) of reduced magmatism in Sierra Nevada–Peninsular Ranges magmatic arc assemblages (Dickinson and Gehrels, 2010a). 6No voluminous igneous assemblages known for North America during this age span (Dickinson and Gehrels, 2009). 7Origins of 16 grains (<0.05% of the total) with ages of 706–888 Ma in nine samples (16% of the total) are indeterminate. 8Three grain ages of 3194–3575 Ma in three separate samples not tabulated. 9Basement of the Archean Wyoming province was masked by Paleozoic–Mesozoic cover during Cretaceous–Paleogene sedimentation.

now considered a basal member of the Green Conglomeratic Intervals mation on the plunging nose of the San Rafael River Formation (Davis et al., 2009b). Colton Swell (Fig. 4). The pebbly beds are interpreted strata include both alluvial and delta-plain For 75–85 km along the southeastern fl ank as a lateral vestige of the Tuscher Formation deposits (Morris et al., 1991), including tan to of the Uinta Basin, basal Colton Formation (Lawton, 1983, 1986b), which was otherwise red channel-form sandstone bodies and inter- includes 10–50 m of pebble conglomerate and removed by erosion across the growing San vening intervals of red to green overbank mud- conglomeratic sandstone of the Paleocene Dark Rafael Swell beneath the unconformity with the stone; the proportions of channel and overbank Canyon sequence (Figs. 3 and 4), a braidplain North Horn Formation (Fig. 4). We infer that deposits are areally variable. Arkosic sandstone succession that unconformably overlies the the clasts in the pebbly beds were derived from is dominant in the core of the Colton sediment Cretaceous (Campanian) Tuscher Formation, Mesozoic strata of the Laramide San Rafael wedge at the Green River. Lithologically similar which is commonly bleached below the con- Swell as it began to grow in the Sevier foreland strata beneath the Green River Formation in the tact (Willis, 1986; Franczyk and Pitman, 1987; before it had developed enough structural relief Piceance Basin (Fig. 1) are termed the DeBeque Franczyk et al., 1990). Sandstone-rich inter- to interrupt sedimentation. Formation (Smith et al., 2008). vals at the top of the Dark Canyon sequence The Colton Formation is exposed continu- locally grade upward into the Colton Formation Green River Formation ously for ~200 km along the face of the Roan (Mathers, 2009). In other places, sharp transi- Cliffs, and extends downdip into the subsurface tions between conglomerate and sandstone Interfi ngering of the Colton Formation with of the Uinta Basin where it grades northward refl ect condensation of section within an evolv- the overlying Green River Formation led to the into the Green River Formation (Fig. 3). Mod- ing fl uvial succession. Dark Canyon clasts are progradation of deltaic bodies as stratigraphic eling Colton sediment volume as a half-cone chert and quartzite pebbles probably derived tongues projecting into lacustrine successions. with its apex at Desolation Canyon of the Green from Mesozoic strata eroded off the growing The Sunnyside delta complex (Remy, 1992; River, where the thickness is ~1000 m (Cashion, Uncompahgre uplift to the southeast (Fig. 1), Schomacker et al., 2010) is a lateral equivalent 1967), yields ~3000 km3 of preserved Colton from which Dark Canyon paleocurrents (Fig. 4) of the uppermost part of the Colton Formation sediment. The original sediment volume was suggest derivation of sediment. as exposed farther east along the Green River greater because no regional thinning of Colton Along the southwestern fl ank of the Uinta (Fig. 4). The Horse Bench Sandstone Bed strata southward toward surface outcrops Basin, lithologically analogous pebbly beds of the Green River Formation (Fig. 4) is the has been detected (Fouch et al., 1976, 1992; (Lawton, 1983, 1986b), 6–18 m thick and com- stratigraphically highest record of delta-related Franczyk et al., 1989). Paleocurrents indicate posed dominantly of chert and quartzite pebbles sediment of arkosic petrofacies (Table 2) in the derivation of Colton sediment consistently from dispersed in pebbly sandstone, are present Uinta Basin, and maintains a consistent thick- southern azimuths, whereas paleocurrent trends unconformably beneath Paleocene strata of the ness of ~6 m for long distances along strike west for underlying units refl ect derivation from North Horn Formation, where they concordantly of the Green River. East of the Green River, the western azimuths (Fig. 4). overlie the Cretaceous (Campanian) Farrer For- Horse Bench Sandstone Bed is <6 m thick, and

858 Geosphere, August 2012

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Ma north of 37°30′ N northeast off the crest of the arch (Fig. 4), sug- south of 37°30′ N 35 Late Duchesne River gesting that Colton and DeBeque depositional systems were separate. DeBeque paleocurrents U i n t a Middle toward the north-northwest in the Piceance 45 e hb n Basin (Lorenz and Nadon, 2002) imply that e Claron c Colton sediment was not transported across the 50 o Green River E Early Douglas Creek arch, although post-Colton Lake “Colton” Colton Uinta eventually overtopped the arch. De ? Beque e Flagstaff n Late “Colton” Formation

60 c Middle Pine Hollow o N o r t h

l In the San Pitch Mountains (Gunnison Pla- a Early H o r n P teau) and Cedar Hills ~75 km southwest of the Canaan Uinta Basin (Fig. 3), Paleogene ﬂ uviatile sand- Maastrichtian Pe a k 70 stone and associated ﬂ oodplain mudstone and siltstone intervening between lacustrine strata of the Flagstaff Limestone Member below and the Late P r i c e Tuscher main body of the Green River Formation above n R i v e r

a Farrer reach thicknesses of 165–245 m, and have been 75 i Kaiparowits

n mapped as the Colton Formation (Marcantiel

a Bluecastle

p and Weiss, 1968; Witkind et al., 1987; Witkind Neslen Grand Capping m Castle Wahweap and Weiss, 1991). The exposures are not con- a Middle Castlegate

C tiguous, however, with the Colton Formation of Wahweap Masuk 80 Blackhawk (l, m, u) and the Uinta Basin, and the quartzose sandstones a n d Muley Early S t a r P o i n t of the succession are unlike the Colton arkosic Canyon Santonian petrofacies (Table 2). In the absence of an alter-

s B l u e

o Gate Straight Cliffs nate name for the so-called Colton Formation of Coniacian c

n the San Pitch Mountains, we refer to the strata

90 a Ferron as “Colton” Formation because they have been Turonian M Tropic Tununk mapped to date only as Colton Formation, but in our view, merit a separate stratigraphic name. Cenomanian Dakota Dakota

100 SEVIER AND MOGOLLON Mussentuchit PROVENANCES Albian a n d n 75–90

i 110 r Ruby Ranch % of

a 45–60 t arc-derived

d For provenance analysis, 52 of the 57 detrital n DZ grains

e 15–35 u Poison Strip zircon samples of Table 1 are grouped into 13

o 120 Aptian C 5–15 (<285 Ma) M Buckhorn stratigraphic and areal subsets (A–M, Table 4). 0–5 Barremian Samples DR (Dome Rock), COL5 (North Horn), 130 GRF (Flagstaff), COL9 (“Colton”), and 12JL5 Figure 2. Schematic Cretaceous–early Paleogene (pre-Oligocene) chronostratigraphy (vari- (Claron) are treated individually. From K-S ous formations and members) of the Utah foreland (units are plotted from west to east analysis, P > 0.05 for comparisons of samples within columns); l—lower, m—middle, u—upper. Scale is expanded for 70–80 Ma and con- within each subset except for subsets A and L, tracted for older than 100 Ma and younger than 45 Ma. Colored dots denote detrital zircon for which we infer that composite detrital zircon (DZ) samples (Table 1) keyed to U-Pb chronofacies (lower right). Adapted for the time scale populations for multiple samples are neverthe- of Walker and Geissman (2009) after Allen and Johnson (2010), Davis et al. (2009a, 2009b), less more reliable indicators of net provenance Eaton (1990, 1991), Eaton et al. (1987, 2011), Fielding et al. (2010), Franczyk and Pitman than the populations of individual samples. (1991), Goldstrand (1990), Goldstrand and Eaton (2001), Goldstrand et al. (1993), Jinnah Sediment was delivered to the Cretaceous and Roberts (2011), Jinnah et al. (2009), Johnson et al. (2011), Larsen et al. (2010), Lawton Sevier foredeep of the southern Cordilleran and Bradford (2011), Lawton et al. (1993, 2003), Ludvigson et al. (2010), Remy (1992), foreland basin both by transverse paleofl ow Roberts (2007), Roberts et al. (2005), Sames et al. (2010), and Smith et al. (2008). eastward off the Sevier thrust belt and by longitudinal paleofl ow northward from the Mogollon highlands transecting central Arizona (Lawton is mapped as the base of the Evacuation Creek Green River Formation are present on the crest et al., 2003; Dickinson and Gehrels, 2008, Member overlying the Parachute Creek Mem- of the arch as lateral equivalents or younger 2010a, 2010b; Lawton and Bradford, 2011). ber (Cashion, 1967). analogues of the Colton Formation to the west The Sevier and Mogollon provenances yielded East of the Green River, the Colton Forma- in the Uinta Basin and the DeBeque Formation contrasting pre-arc detrital zircon populations. tion thins as it onlaps the fl ank of the Laramide to the east in the Piceance Basin (Fig. 3). Paleo- Detritus from the Mogollon provenance (Fig. Douglas Creek arch (Fig. 4). Only thin sandstone currents in the basal Paleogene sandstones on the 5D) was dominated by subpopulations IV intervals mapped either below or within the basal Douglas Creek arch diverge to the northwest and (anorogenic granite) and V (Yavapai-Mazatzal)

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P r e c a m b r i a n c o r e o f U i n t a u p l i f t 111° W 109° W WY 50 km

Neogene cover of N Browns Park Fm

Sevier CCF U i n thrust t a n f r o n t e e D r r G e R v 40° N Ri UT CO Me B a 40° N IND CO22 s Piceance GRF i n CHC COL5 COL3 DCA basin WRU N GSD COL4 CP34 WADC COL2 P COL1 COL6 DC1 DEB r COL9 e v SPC COL8 i 06-8B R COL7 GC S GJ DZ samples 109° W CP33 S a n GR DCS towns Rafael RRR54 o d Post - Colton Eocene a S w e l l CP32 r Laramide basin fill o (Green River, Uinta, RRR12 MVF ol Dushesne River Fms) C

111° W Mo Paleocene - Eocene Colton Formation and correlatives (DeBeque, "Colton" Fms) pre - Upper pre - Maastrichtian Maastrichian - Paleocene Cretaceous Upper Cretaceous North Horn - Flagstaff interval

Figure 3. Geologic sketch map of the Laramide Uinta Basin and nearby areas adapted after Cashion (1973), Gualtieri (1988), Hintze (1980), Tweto (1979), Jefferson (1982), Weiss et al. (1990), Witkind (1988, 1995), Witkind and Weiss (1991), and Witkind et al. (1987). For reasons of scale, stratigraphic units (Fig. 2) are not shown where <100 m thick, contacts are smoothed in areas of intricate topography, isolated exposures are not plotted in areas of complex structure, and local Neogene cover is omitted. The subsurface extent of the Colton Formation is adapted after Fouch et al. (1976), Fouch (1981), Franczyk et al. (1989), Franczyk (1991), and Sprinkel (1994). See Table 1 for 29 detrital zircon (DZ) samples, except that GC denotes 7 DZ samples from Gray Canyon (Table 1) of the Green River, and 4 samples (TF1 and DC2, DC3, DC4) of Mathers (2009) were collected near locality COL7 in nearby Tusher Canyon (tributary to the Green River). Key geologic features: CCF—Cretaceous Currant Creek Formation (Isby and Picard, 1983), DEB—DeBeque Formation (Colton-equivalent) on the south ﬂ ank of the Piceance basin, DCA—Douglas Creek arch, DCS—lateral extent of the conglomeratic Dark Canyon sequence at the base of the Colton Formation east of the Green River (after Franczyk and Pitman, 1987), SPC—“Colton” Formation (i.e., Colton correlative) of the San Pitch Mountains (CHC denotes related exposures in the Cedar Hills), GSD—Sunnyside delta and overlying Horse Bench Sandstone Bed in the Green River Formation at Gate Canyon, IND—Upper Cretaceous Indianola Group near the Sevier thrust front, MVF—Oligocene Marysvale volcanic ﬁ eld, WRU—White River uplift. States: CO—Colorado, UT—Utah, WY—Wyoming. Towns: D—Duchesne, GJ— Grand Junction, GR—Green River, Me—Meeker, Mo—Moab, N—Nephi, P—Price, R—Rangely, S—Salina, V—Vernal.

of Table 3, whereas detritus from the Sevier Subordinate proportions of subpopulations reduced proportion of subpopulations IV and V provenance (Fig. 5ABC) was dominated by IV and V are present in Sevier detritus because for the southern Sevier provenance (Fig 5C), subpopulations II (Paleozoic–Neoproterozoic) those subpopulations are present in stratigraphic as opposed to the northern Sevier provenance and III (Grenville) recycled from Paleozoic and units of the thrust belt and were recycled (Fig. 5AB), is interpreted to reﬂ ect derivation Mesozoic strata of the thrust belt (Dickinson together with subpopulations II and III (Dickin- from a different mix of strata within the thrust and Gehrels, 2008; Larsen et al., 2010). son and Gehrels, 2008; Lawton et al., 2010). The belt. The age range of recycled detrital zircons

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Figure 4. Stratigraphic tran- W Bench E o r s e sect of middle Campanian to H COL3 S a n Green River Fm d s t o n e Middle Eocene strata along the Bed G r e (ca. 4 south ﬂ ank of the Uinta Basin Sunnyside e n 7.5 M Ri a) from the Wasatch Plateau delta v e r complex COL4 (west, W) to the Douglas Creek Fm arch (east, E). Colton includes the Renegade Tongue of the Green COLTON Fm Green River Formation where River Dark Canyon Fm COL7 sequence the two are nearly amalga- COL5 (10–50 m) COL1 COL22 DC2- DC1 mated between the Green River DC4 DC3 COL8 Flagstaff pebbly beds and Westwater Canyon. Paleo- Limestone (6–18 m) GRF Fm current trends are azimuthal Tuscher Douglas COL5 TF1 vectors with north to the top. MT9 North MT1 Creek See Table 1 for detrital zircon Farrer Fm Horn s arch (DZ) samples. Adapted after MF4 MF6 castle S plunging Blue Fms Fm n - Sego Cashion (1967, 1973), Chan CP34 MB10 Nesle Ss tlegate nose of MN7 Cas and Pfaff (1991), Dickinson MN8 e ric m et al. (1986), Fouch et al. (1976, P F San Rafael 500 m er iv s 1983), Franczyk and Pitman R S S w e l l te a scale g (1987, 1991), Franczyk et al. le st a (1990, 1991), Gualtieri (1988), C Wasatch 50 25 fluvial DZ km km Lawton (1983, 1986b), Lawton Plateau paleocurrents samples and Bradford (2011), Mathers (2009), McLaurin and Steel Soldier Willow Soldier Horse Green Tusher Westwater UT/CO Summit Creek Creek Canyon River Canyon Canyon Line (2007), Miall and Arush (2001), Moncure and Surdam (1980), Olsen (1995), Olsen et al. (1995), Peterson (1976), Remy (1992), Robinson and Slingerland (1998), Ryder et al. (1976), Schomacker et al. (2010), Van de Graaf (1972), Willis (2000), Witkind (1988, 1995), Weiss et al. (1990), and Zawiskie et al. (1982). CO—Colorado, UT—Utah.

varies longitudinally in the Cordilleran foreland (100 × subpopulation I of Table 3/total detrital recycled from strata uplifted along the Sevier basin because the detrital zircon populations of zircons) indicates the proportion of arc-derived thrust belt. strata incorporated into the thrust belt vary along grains in detrital zircon populations, and ranges For a Mogollon index >60, arc-derived grains strike (Leier and Gehrels, 2011). from 0 to ~75 (an arc index of 100 would indi- were by inference added to pre-arc detrital zir- The paucity of arc-derived detrital zircons cate arc-derived grains only). The Mogollon con populations containing signiﬁ cant propor- in sand spread eastward into the foredeep from index (subpopulations IV + V/total pre-arc tions of detritus from the Mogollon provenance. sources along the Sevier thrust belt implies that grains) indicates the proportion of pre-arc grains Mogollon indices in the range of 60–95 imply arc-derived detritus reaching the Utah foreland derived from the Paleoproterozoic Yavapai- admixtures of detritus from Sevier and Mogol- did not travel transversely across the Sevier oro- Mazatzal belt of southwest Laurentia as intruded lon provenances in varying proportions, with gen from the Sierra Nevada segment of the Cor- by anorogenic granites of Mesoproterozoic age, the maximum Mogollon index of ~95 observed dilleran magmatic arc directly to the west, but and ranges from 20 to ~95 (a Mogollon index of only for subset M (Table 4), the Mogollon prov- instead moved longitudinally along the foredeep 100 would indicate pre-arc grains exclusively of enance signature of Figure 5D, and sample DR east of the thrust front from segments of the arc subpopulations IV and V). (Dome Rock Upper Cretaceous of Table 1) farther south beyond the southern limit of Sevier A plot of arc index against Mogollon index collected south of the limit of Sevier thrust- thrusting (Fig. 1). Intramontane Cretaceous– indicates the relative contributions of detrital ing (Fig. 1). The Mogollon index of subset B Paleogene deposits within the Sevier orogen zircon grains from the two provenances in rela- (Colton) is intermediate between Mogollon between the Sierra Nevada and the Sevier thrust tion to contributions from the magmatic arc indices for Sevier and Mogollon provenances belt contain few arc-derived detrital zircon (Fig. 6). A Mogollon index of 60 delimits the (Fig. 6), suggesting that detritus from both prov- grains (Druschke et al., 2011). Minor detrital ﬁ eld of Sevier-derived detritus with arc indi- enances reached the California paleoriver carry- zircon grains of Jurassic and Cretaceous age ces <10. For samples with a Mogollon index ing Colton sediment (Fig. 1). may have reached the foredeep from backarc <60, the plot suggests that arc-derived grains For the Utah foreland, arc indices >30 are igneous centers (du Bray, 2007) located east of were added to pre-arc detrital zircon popula- observed only for strata of late Campanian and the main arc trend along the Sierra Nevada. tions derived largely if not exclusively from the younger age and Mogollon indices >60 are Sevier provenance. The ancestry of individual observed only for Paleogene and the youngest Grain Age Indices grains of subpopulations IV and V is equivo- preserved late Campanian (Tuscher) strata (Fig. cal, however, because grains derived directly 2). These relations suggest coupled enhance- For analysis of Cretaceous–Paleogene depo- from the Mogollon highlands cannot be distin- ment over time of both arc and Mogollon detri- sitional patterns in the backarc foreland of Utah, guished in derivative sandstones from grains tus transported longitudinally from south to we employ two grain age indices. The arc index derived initially from southwest Laurentia but north within the foreland region.

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TABLE 4. P VALUES (INTERSAMPLE COMPARISONS) FROM K-S ANALYSIS OF U-Pb AGES OF DETRITAL ZIRCONS IN SUBSETS OF CRETACEOUS–PALEOGENE SAMPLES FROM THE SEVIER FORELAND OF UTAH AND THE FLANK OF THE ARIZONA–NEW MEXICO MOGOLLON HIGHLANDS Arc-derived DZ Samples Total DZ grains (younger Pre-arc DZ grains Stratal subsets (see Table 1) grains than 285 Ma) (older than 285 Ma) Uinta Basin and Table Cliffs syncline (post–middle Campanian) A—Sunnyside delta1 (N = 2) COL3, COL4 0.01 nd2 0.02 B—Colton Formation3 (N = 5) CO22, COL1, 6, 7, 8 0.11–0.99 0.59–1.00 0.11–0.99 B′—Colton Formation3 (N = 4) DC1, DC2, DC3, DC4 0.17–0.87 0.19–0.89 0.52–0.99 B + B′—Colton Formation3 (N = 9) See B and B′ NA4 0.15–0.99 0.10–0.99 C—DeBeque Formation5 (N = 2) WADC, COL2 0.27 nd2 0.74 D—Tuscher Formation (N = 2) MT9, MT1 0.54 0.05 0.95 D′—Tuscher Formation (N = 3) MT9, MT1, TF1 NA4 0.05–0.99 0.90–0.96 E—Farrer Formation (N = 2) MF4, MF6 0.58 0.99 0.08 F—Kaiparowits6 (N = 5) 4JL5, KK2, 4, 5, 6 0.00–1.007 0.06–0.97 0.50–1.00 G—Maastrichtian–Paleocene of Table Cliff basin (N = 6) 11JL5, 9JL5, 8JL5, 7JL5, 6JL5, CP18 0.10–0.99 nd2 0.10–0.98 Sevier foredeep (pre–late Campanian) and fl ank of Mogollon Highlands H—Upper Cretaceous of northern9 Sevier provenance (N = 4) MB10, MN7, MN8, CP34 0.09–0.84 nd2 0.37–0.90 I—Lower Cretaceous of northern9 Sevier provenance (N = 3) 5-6B, CP32, RRR12 0.04–0.6010 nd2 0.04–0.7910 J—Upper Cretaceous of southern9 Sevier provenance (N = 5) P6GC, WF6GC, WP9GC, 2JL5, CP4011 0.19–0.98 nd2 0.48–0.99 K—Upper Cretaceous of mixed Sevier–Mogollon provenance KK1, CP3312, CP39, COL1112, 1JL513 0.00–0.997 0.0614 0.07–0.85 (longitudinal paleofl ow) (N = 5) L—mid-Cretaceous (Dakota) of uncertain provenance (N = 2) COL10, RRR54 0.03 0.00 0.07 M—Cretaceous–Paleogene of Mogollon provenance (N = 4) NARCHU, COL12, COL13, CP9 0.00–0.3315 nd2 0.00–0.5015 Note: N indicates the number of samples per subset. P—probability values calculated from Kolmogorov-Smirnoff (K-S) statistics. DZ—detrital zircon. Bold type denotes P ≥ 0.05. Five samples (12JL5, COL9, GRF, COL5, DR) of Table 1 are not included in any tabulated subsets. 1Lower Green River Formation of Sunnyside delta complex in Gate Canyon (laterally equivalent to upper Colton Formation) and younger Horse Canyon Sandstone Bed (Fig. 4). 2Not determined (K-S analyses are unreliable for <20 U-Pb ages per DZ subpopulation). 3Colton Formation, including the basal Dark Canyon sequence (conglomeratic). 4Not calculated (nonrandom selection of grains for dating in B′ samples and sample TF1). 5Colton-equivalent from the Piceance Basin of northwestern Colorado. 6Middle and upper Kaiparowits Formation (see subset K for lower Kaiparowits Formation). 7Large range in P values from variable proportions of arc and pre-arc DZ subpopulations (sample 4JL5 untabulated). 8Sample CP1 not used for calculation of P for total DZ grains (nonrandom selection of grains for dating). 9Dividing line between northern and southern Sevier provenances at 47°30′ N for foredeep samples from stratigraphic units with paleocurrents indicating eastward transverse paleoflow off the Sevier thrust belt. 10P < 0.05 solely from the CP32–RRR12 sample comparison. 11Anomalous sample CP40 yields untabulated P = 0.01–0.13 for all grains and P = 0.02–0.11 for pre-arc grains. 12P = 0.92 for total grains and P = 0.83 for pre-arc grains in Last Chance (CP33) and Notom (COL11) delta lobes (Garrison and van den Burgh, 2004) of Ferron delta complex. 13Not used for total DZ grain population. 14For the KK1–1JL5 sample comparison only (<20 arc grains per sample in other samples). 15Large range in P values stems from variable proportions of detritus derived from closely associated anorogenic granite and Yavapai-Mazatzal sources (DZ subpopulations IV and V of Table 3) of Mogollon provenance.

I II III IV V 95 A 160 415 1775 subset H – northern Sevier 590 1115 1520 Upper Cretaceous Figure 5. Age distribution N = 4 n = 382 curves of composite detrital zir- 150 425 1110 con (DZ) samples illustrative of B subset I – northern Sevier 590 1695 detritus from the Sevier thrust 1500 Lower Cretaceous belt (A–C) and the Mogollon N = 3 n = 288, highlands (D). Selected peak 415 80 1095 ages are labeled to the nearest C 155 595 subset J – southern Sevier Upper Cretaceous 5 m.y. Colored bands (I–V) are 1740 the subpopulations of Table 3. N = 5 n = 403 See Table 4 for sample subsets. 1700 Arc-derived (younger than 285 Ma) subpopulation I forms only 3%–5% of the detrital zircons in the DZ populations plotted. N is the number of D subset M – Arizona Mogollon samples composited, and n is Upper Cretaceous and Paleogene N = 4 n = 382 the number of dated grains in 1405 each composited population. The division between north- 160 ern and southern Sevier prov- 100 ′ enance is at 47°30 N (Fig. 2). 1120

0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 3250 Age (Ma)

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Figure 6. Plot of arc index however, that P < 0.05 for comparisons of 80 (ordinate) versus Mogollon arc + Mogollon Paleogene Colton of the Uinta Basin (subset B) index (abscissa) for detrital (± minor Sevier) with Paleogene DeBeque of the Piceance Basin F – Kaiparowits zircon (DZ) populations in DZ (subset C), and with most Cretaceous subsets sample subsets of Table 4 (solid 60 arc (D–F, K, L). plus dots) and other individual Fla For pre-arc grains in most samples of Tables Sevier DR

) samples of Table 1 (open dots): I DZ B – Colton 5 and 6, P > 0.05 for 34% of sample pairs

( E – Farrer Cla—Claron (12JL5), “Cot”— 0 6 (Table 7), more than for total grain populations

e 40

= “Colton” (COL9), NH—North d I (10%) but fewer than for arc-derived grains

c Horn (COL5), Fla—Flagstaff r D – Tuscher (73%) in the same samples. Arc-derived grains

a Sevier (GRF), DR—Dome Rock. provenance L – Dakota having similar age spectra were evidently added C – DeBeque Index numbers for subsets are 20 G – Table to pre-arc grains of varied Sevier or Mogollon or Cliff K – mixed mean values. Sevier (recycled) foredeep mixed provenance. Note that P > 0.05 for com- A – Sunnyside provenance subsets: H—north- provenance parisons of Colton (subset B) with the underly- I M – Mogollon ern Sevier Upper Cretaceous, J Cla NH H “Cot” ing Tuscher Formation (subset D) and for the 0 I—northern Sevier Lower Cre- 020406080100Farrer Formation (subset E), which underlies taceous, J—southern Sevier Mogollon index (MI) the Tuscher Formation, with the laterally equiv- Upper Cretaceous. alent Kaiparowits Formation (subset F) of the Table Cliff basin. Conversely, P < 0.05 for either Colton or Tuscher grains in comparison with STATISTICAL COMPARISONS and of subset F (Campanian Kaiparowits) with either Farrer or Kaiparowits grains. These rela- sample GRF (Paleogene Flagstaff underlying tions suggest progressive provenance evolution K-S comparisons of total detrital zircon pop- Colton). Separate K-S analyses of arc-derived in the Utah foreland over the interval spanning ulations in the sample subsets of Table 4 and and pre-arc subpopulations help to clarify prov- Kaiparowits–Farrer and Tuscher–Colton For- selected individual samples (Table 1) having an enance relations. mations deposition (Fig. 2). arc index >10 show that P < 0.05 for 93% of For arc-derived grains considered separately K-S comparisons for pre-arc grains in sam- sample pairs (Table 5), largely because variable from pre-arc grains in the samples of Table 5, ples with contrasting Mogollon indices (>60 proportions of arc-derived and pre-arc grains P > 0.05 for 78% of sample pairs (Table 6), and <60) consistently yield P ≤0.05 and are not reduce P values. P is high (>0.25) only for com- suggesting that age spectra of detrital zircons tabulated. For pre-arc grains in sample subsets parisons of subset B (Paleogene Colton) with in arc detritus delivered to the backarc Utah and individual samples of varied Cretaceous– sample DR (Upper Cretaceous from the Mojave foreland were largely comparable during Late Paleogene ages but with Mogollon index <60, region), a relation noted in Davis et al. (2010), Cretaceous and early Paleogene time. Note, P > 0.05 for only 38% of sample pairs (Table 8),

TABLE 5. MATRIX OF COMPARATIVE P VALUES FROM K-S ANALYSIS OF TOTAL DETRITAL ZIRCON POPULATIONS IN SAMPLE SUBSETS AND INDIVIDUAL SAMPLES (GRF, DR) HAVING AN ARC INDEX >10 A B C GRF D E F G K L DR A—Sunnyside delta x – 0.07 –0.04–––––– B—Colton Formation – x – – – 0.02 ––––0.38 C—DeBeque Formation 0.07 –x –0.23 –––0.01–– GRF—Flagstaff Limestone – – – x – – 0.57 –––0.02 D—Tuscher Formation 0.04 – 0.23 – x–––––– E—Farrer Formation – 0.02 – – – x ––––0.02 F—Kaiparowits Formation – – – 0.57 ––x–––– G—Table Cliffs strata – – – – – – x – – – K—Upper Cretaceous foredeep – – 0.01 – – – 0.01 x – – L—Dakota Formation – – – –––––x– DR—Dome Rock succession – 0.38 –0.02 ––––x Note: P—probability values calculated from Kolmogorov-Smirnoff (K-S) statistics. Bold type denotes P > 0.05. P values are rounded to the nearest 0.01. Dashes indicate P = 0.00. For sample subsets, see Table 4. For samples having an arc index >10, see Figure 6.

TABLE 6. MATRIX OF COMPARATIVE P VALUES FROM K-S ANALYSIS OF ARC-DERIVED DETRITAL ZIRCON SUBPOPULATIONS (U-Pb AGE YOUNGER THAN 285 MA) IN SAMPLE SUBSETS AND INDIVIDUAL SAMPLES (GRF, DR) HAVING AN ARC INDEX >10 A B C GRF D E F G K L DR A—Sunnyside delta x 0.20 0.24 0.49 0.48 0.18 0.76 0.81 0.64 0.06 0.15 B—Colton Formation 0.20 x–0.87 0.01 – – 0.78 0.02 0.02 0.21 C—DeBeque Formation 0.24 –x0.14 0.99 0.32 0.45 0.18 0.29 0.03 0.04 GRF—Flagstaff Limestone 0.49 0.87 0.14 x 0.31 0.05 0.48 0.64 0.39 0.40 0.82 D—Tuscher Formation 0.48 0.01 0.99 0.31 x 0.30 0.55 0.39 0.34 0.04 0.06 E—Farrer Formation 0.18 – 0.32 0.05 0.30 x 0.30 0.13 0.30 0.38 0.01 F—Kaiparowits Formation 0.76 – 0.45 0.48 0.55 0.30 x 0.10 0.41 0.03 0.08 G—Table Cliffs strata 0.81 0.78 0.18 0.64 0.39 0.13 0.10 x 0.48 0.35 0.27 K—Upper Cretaceous foredeep 0.64 0.02 0.29 0.39 0.34 0.30 0.41 0.48 x 0.50 0.16 L—Dakota Formation 0.06 0.02 0.03 0.40 0.04 0.38 0.03 0.35 0.50 x 0.19 DR—Dome Rock succession 0.15 0.21 0.04 0.82 0.06 0.01 0.08 0.27 0.16 0.19 x Note: Bold type denotes P ≥ 0.05. P—probability values calculated from Kolmogorov-Smirnoff (K-S) statistics. P values are rounded to the nearest 0.01. Dashes indicate P = 0.00. For sample subsets, see Table 4. For samples having an arc index >10, see Figure 6.

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TABLE 7. MATRIX OF COMPARATIVE P VALUES FROM K-S ANALYSIS OF PRE-ARC DETRITAL ZIRCON SUBPOPULATIONS (OLDER THAN 285 MA BY U-Pb AGE) IN SAMPLE SUBSETS AND SAMPLE DR HAVING AN ARC INDEX >10 AND A MOGOLLON INDEX >60 OR A MOGOLLON INDEX <60 BUT AN ARC INDEX >30 AB C DEFDR A—Sunnyside delta x 0.01 0.68 0.11 ––0.10 B—Colton Formation 0.01 x 0.01 0.58 – 0.01 0.04 C—DeBeque Formation 0.68 0.01 x 0.10 – – 0.03 D—Tuscher Formation 0.11 0.58 0.10 x––0.04 E—Farrer Formation – – – – x 0.36 – F—Kaiparowits Formation – – – – 0.36 x– DR—Dome Rock succession 0.10 0.04 0.03 0.04 – – x Note: P—probability values calculated from Kolmogorov-Smirnoff (K-S) statistics. Bold type denotes P ≥ 0.05. P values are rounded to the nearest 0.01. Dashes indicate P = 0.00. GRF of Tables 5 and 6 omitted (<20 pre-arc grains, too few for reliable K-S analysis). See Table 8 for sample subsets G, K, and L of Tables 5 and 6 (P = 0.00 for K-S comparisons of Table 8 samples with samples tabulated here). For arc index and Mogollon index data, see Figure 6. For sample subsets, see Table 4; for sample DR, see Table 1.

TABLE 8. MATRIX OF COMPARATIVE P VALUES FROM K-S ANALYSIS OF PRE-ARC DETRITAL ZIRCON SUBPOPULATIONS (OLDER THAN 285 MA BY U-Pb AGE) IN SAMPLE SUBSETS AND INDIVIDUAL SAMPLES (12JL5, COL5, COL9) HAVING A MOGOLLON INDEX <60 12JL5 COL9 COL5 E F G H I J K L 12JL5—Claron x – – –––––0.22 –0.04 COL9—“Colton” – x 0.81 0.22 0.66 – 0.14 ––0.09 – COL5—North Horn – 0.81 x 0.19 0.91 – 0.52 0.01 – 0.01 – E—Farrer – 0.22 0.19 x 0.36 – 0.34 0.10 – 0.98 0.02 F—Kaiparowits – 0.66 0.91 0.36 x–1.00 0.01 – 0.09 – G—Table Cliff syncline – – – – – x – 0.13 ––0.66 H—northern Sevier-UK – 0.14 0.52 0.34 1.00 – x –––– I—northern Sevier-LK – – 0.01 0.10 0.01 0.13 –x–0.16 0.29 J—southern Sevier-UK 0.22 – – –––––x–0.17 K—mixed foredeep – 0.09 0.01 0.98 0.09 ––0.16 –x0.01 L—Dakota (foredeep) 0.04 – – 0.02 – 0.66 – 0.29 0.17 0.01 x Note: P—probability values calculated from Kolmogorov-Smirnoff (K-S) statistics. LK—Lower Cretaceous; UK—Upper Cretaceous. Bold type denotes P ≥ 0.05. P values are rounded to the nearest 0.01. Dashes indicate P = 0.00. Farrer (E) and Kaiparowits (F) repeated from Table 7. For sample subsets, see Table 4. For samples having a Mogollon index <60, see Figure 6.

suggesting mixing in varying proportions of grains in both samples are similar to those for Upper Campanian strata (Fig. 9BCD) in detritus from Sevier and Mogollon provenances sandstones derived mainly from the northern the Utah foreland contain in net as many arc- in varying patterns over time within the Utah Sevier thrust belt (Figs. 5D and 8C), with P = derived grains (~50%) as the Colton Formation foreland. 0.14 (“Colton”) and P = 0.52 (North Horn) from (Fig. 9A), and distinctly more than older Upper K-S comparisons of pre-arc grains (Table 8). Cretaceous strata (Fig. 9E). Paleocurrents and UINTA BASIN PROVENANCE Facies relations in the San Pitch Mountains sug- facies patterns in the pre-upper Campanian gest that “Colton” sand was derived exclusively strata refl ect components of longitudinal paleo- Detrital zircon populations from Paleogene from nearby frontal Sevier thrust sheets (Mar- fl ow from the southwest in the Utah foreland strata of the Uinta Basin and nearby areas (Figs. cantiel and Weiss, 1968), and North Horn paleo- (am Ende, 1991; Lawton et al., 2003; Garrison 3 and 4) fall into three groups refl ecting differ- currents (Fig. 4) are compatible with derivation and van den Burgh, 2004; Janok et al., 2010), ent provenance relations (Fig. 7): (1) samples from the Sevier thrust belt to the west. even though the net proportion of arc-derived from the North Horn and “Colton” Formations grains is much lower than in upper Cam panian that contain ≤1% arc-derived grains, (2) sam- Colton Petrofacies units. These relations suggest that some arc- ples from the Colton Formation, including the derived detritus began to reach the backarc Dark Canyon sequence, that contain 44%–58% The Paleogene Colton Formation and Creta- region of Utah with the initiation of the Cor- arc-derived grains, and (3) samples from the ceous strata of the Utah foreland contain simi- dilleran foreland basin in mid-Cretaceous time, DeBeque Formation and Sunnyside delta com- lar arc-derived detrital zircon subpopulations but not in abundance until Campanian time, after plex, including the Horse Bench Sandstone Bed, (Fig. 9), although proportions of arc-derived which delivery of abundant arc detritus to the that contain an intermediate level of 12%–23% grains are variable. All three arc-derived subpopu- Utah foreland continued into Paleogene time. arc-derived grains. lations (Ia–Ic of Table 3) are present. Samples Fluvial transport of arc-derived grains into the from the conglomeratic Dark Canyon sequence Uinta Basin from the south by longitudinal Sevier Provenance (basal Colton Formation) were composited rather than transverse paleofl ow with respect with other Colton samples because our single to the Sevier thrust belt to the west is inferred The Maastrichtian–Paleocene North Horn sample from the Dark Canyon sequence (COL8 because units containing abundant arc-derived Formation (COL5 of Table 1) exposed along of Table 1) yields P = 0.39–1.00 for all grains, zircons (>25%) display northeastward to north- the southwest fl ank of the Uinta Basin and the P = 0.85–1.00 for arc-derived grains, and P = westward paleocurrents, whereas units lacking Paleocene–Eocene “Colton” Formation (COL9 0.51–0.99 for pre-arc grains when compared by many arc-derived zircons (<10%) display east- of Table 1) in the San Pitch Mountains south- K-S analysis to our other four Colton samples, ward to southeastward paleocurrents refl ective west of the Uinta Basin (Figs. 2 and 3) contain and the single Mathers (2009) Colton sample of sediment transport off the Sevier thrust belt essentially no arc-derived detrital zircons (Fig. from above the Dark Canyon sequence (DC3 of (Fig. 10). 8AB), and P = 0.81 from K-S comparison of Table 1) yields P = 0.31–0.99 from comparison No strata of Maastrichtian and Early Paleo- their full age spectra. The age spectra of pre-arc with her three Dark Canyon samples. cene age in the Uinta Basin (Fig. 2) contain

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COL 2 COL 3 (12%) DeBeque (23%)

)

0 500 1000 1500 2000 i

R s 0 500 1000 1500 2000

S COL 4 (16%) ( COL 8 Dark Canyon (52%)

0 500 1000 1500 2000 0 500 1000 1500 2000

St Ub

"Colton" (0%)

0 500 1000 1500 2000 N COL 7 (44%) O North Horn (1%) T L O 0 500 1000 1500 2000 C 0 500 1000 1500 2000

COL 6 (58%) COL 1 (55%)

0 500 1000 1500 2000 0 500 1000 1500 2000

Figure 7. Summary age distribution curves (abscissas in Ma) for detrital zircon (DZ) populations younger than 2250 Ma (each plot repre- sents 83–94 grain ages) in samples from the Colton Formation and related Paleogene strata (Eocene and Upper Paleocene stippled on inset map) of the Uinta Basin and nearby areas (Fig. 3). Detrital zircon grains older than 2250 Ma (<5% of the total) not plotted for reasons of scale. Pink bands denote arc-derived subpopulation I (see Table 3) where >1% present (percentages in parentheses). See Figure 4 for stratigraphic units (Dark Canyon is basal Colton east of the Green River). St—Sevier thrust belt, Pb—Piceance Basin, Ub—Uinta Basin.

abundant arc-derived zircons. The Upper feldspathic sandstones as the proportion of The Mogollon index (Fig. 7) is <60 for the Paleocene to Lower Eocene Colton Formation arc detritus reaching the basin increased dur- lower three curves but >60 for the upper two (Fig. 2) either unconformably overlies subjacent ing late Campanian (Farrer and Tuscher) and curves, which also reﬂ ect reduced contribu- Campanian strata or is separated from the Cam- Paleogene (Colton and Green River) time. tions of detrital zircon subpopulations II and panian strata by the intervening Sevier-derived North Horn sedimentation reﬂ ected a tran- III recycled from the Sevier thrust belt. For sediment wedge of the Maastrichtian to Lower sient return to the Sevier-derived quartzolithic Tuscher and Colton subsets (Figs. 12A, 12B) of Paleocene North Horn Formation (Fig. 4). There petrofacies in Maastrichtian to Early Paleo- most arkosic composition (Fig. 11) and highest is no internal evidence within the Uinta Basin cene time. Cretaceous Indianola and Paleogene Mogollon index (Fig. 7), feldspar may well have for the persistence of longitudinal foreland “Colton” sandstones deposited near the thrust been derived from Precambrian basement of the transport of arc detritus between late Campanian front contain petrofacies intermediate between Mogollon provenance as well as from plutons and Late Paleocene time during peak Laramide Sevier-derived quartzolithic and Colton arkosic or volcanic assemblages of the Cordilleran arc deformation. petrofacies. in the Mojave region. The close similarity of The Colton arkosic petrofacies represented The age spectra of pre-arc detrital zircon Kaiparowits (Fig. 12D) and Farrer (Fig. 12C) the culmination of a persistent trend of com- grains in Colton and older sandstones of the age spectra for pre-arc grains, and their con- positional evolution for sandstones of the Utah foreland reinforce the interpretation that trast with mutually similar Tuscher (Fig. 12B) Uinta Basin (Fig. 11). Quartzolithic sand- longitudinal sediment transport along the trend and Colton (Fig. 12A) age spectra, parallel the stones derived from the Sevier thrust belt of the foreland basin increased in importance results obtained from K-S comparisons of those were progressively supplanted by increasingly over Cretaceous–Paleogene time (Fig. 12). sample subsets (Table 7).

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III IV V 1055 –1115 1450 1700 A. COL9 : "Colton" Formation (Colton- 1850 equivalent) of San Pitch Mountains [Upper Paleocene - Lower Eocene] N=1 n=93

2660 Figure 8. Comparative age distribution curves for “Colton,” III 1880 1035–1160 North Horn, and northern 1760 Sevier (subset H of Table 4) B. COL5 : North Horn Formation detrital zircon (DZ) popula- 550 of Uinta basin (Willow Creek) tions. Color bands denote sub- 240 1455 [Maastrichtian - Lower Paleocene] populations I–V (Table 3). N is N= 1 n=91 the number of samples (>1 for 2615 -2760 bottom curve only), and n is the number of dated grains in each DZ population plotted. 415 1775 C. Castlegate-Neslen-Bluecastle (Campanian Upper Cretaceous 1115 590 1520 of Book Cliffs and Gray Canyon) N = 4 (subset H) n = 382

2750

0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 3250 Age (Ma)

Figure 9. Comparative age distribution 1a 1b 1c 168 curves for arc-derived (younger than ARC DZ GRAINS <285 Ma A Colton Formation Roan Cliffs 285 Ma) detrital zircon (DZ) subpopula- (Upper Paleocene- tions in Colton (A) and other sample subsets 94 Lower Eocene) 85 N=5 n=221 (B–E) from the Utah foreland (see Table 4) (52%) with arc index ≥15. Color bands denote subset B 76 subpopulations Ia, Ib, and Ic (see Table 3). Figures on the left denote the percentages of 7082 95 156 B Tuscher Formation arc-derived grains in the total grain popula- Gray Canyon (upper Campanian tion of each sample subset. N is the number Cretaceous) of samples composited for each age spec- (28%) N=2 n=52 trum, and n is the number of dated grains in subset D each composite DZ population. 152

98 C Farrer Formation 84 Gray Canyon Dark Canyon Sequence (upper Campanian 76 Cretaceous) Detrital zircon populations in the conglom- (48%) N=2 n=92 eratic Dark Canyon sequence at the base of subset E 81 76 173 the Colton Formation and within the overlying 180 D Kaiparowits Formation Kaiparowits Plateau main body of the Colton Formation are simi- 92 149 lar (Table 4), despite contrasts in comparative 98 (upper Campanian (74%) Cretaceous) lithology. The gravel fraction of the Dark Can- subset F N=5 n=330 yon sequence is a mature pebble lag apparently 96 reworked from underlying upper Campanian 160 Tuscher Formation beneath the basal Colton E pre-upper Campanian Upper Cretaceous unconformity (Fig. 4). The sand fraction of the (Wahweap-Ferron-Dakota) Dark Canyon sequence is less feldspathic and (15%) Utah foreland subsets N=6 n=107 more lithic than the Colton arkosic petrofacies K and L (Fig. 11), with a signiﬁ cant proportion of chert 0 50 100 150 200 250 300 and quartzite sand grains (Table 2) presumably Age (Ma)

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related to the pebble fraction of the Dark Can- A B C D Stratal Wasatch Plateau Roan Cliffs and Desolation and Dark Canyon yon sequence. units including Price Book Cliffs Gray Canyons sequence east Reworking of Colton sand from the Tuscher River canyon (between A and C) of Green River of Green River Formation is not favored for several reasons: 2° 351° 341° 316° (1) the arc-derived detrital zircon subpopulation (I of Table 3) is more abundant by a factor R=0.98 Colton of nearly two in Colton sand as compared to R= N=4 N=7 0.92 N=4 Tuscher sand (44%–58% vs. 28%–31%), and n=220 N=12 R=0.94n-305 R=0.79 n=90 n=80 zircons could not be sorted so strongly by age during recycling; (2) Colton arkosic sandstone is R=0.69 R=0.64 consistently more feldspathic than Tuscher sand- thin 71° stone (Fig. 11), whereas recycling is expected to North feather Horn edge enhance, rather than reduce, the ratio of quartz 87° KEY N=16 N=13 only to feldspar; and (3) the Colton Formation is ﬁ ve n=695 n=530 percentage of times as thick as the Tuscher Formation (Fig. 4), arc-derived DZ grains which in any case was masked from erosion 42° N=2 R=0.98 (<285 Ma) within the Uinta Basin once the Dark Canyon n=60 80° sequence was deposited above it. Tuscher r To explain the close similarity of the detrital e 49° 45%–55%

v N=5

i zircon populations in the basal Dark Canyon R=0.93 n=205 R R=0.98 sequence and the remainder of the Colton For- mation, we suggest a sedimentological rationale

e 25%–45%

c N=5 N=5 58° by which locally derived chert-rich and pebbly i N=3 70°

r n=855 n=145 n=275 detritus was mixed on a conglomeratic braid- P Farrer plain with arkosic sand of distal origin during <10% R=0.99 the earliest phase of Colton sedimentation. Inti- R=0.93 mate intercalation of lenses of cross-bedded sandstone and massive pebble conglomerate N=4 N=10 64° n=510 70° n=165 N=4 ~0% within the Dark Canyon sequence are com- Bluecastle n=135 patible with that viewpoint. The similarity of ( - Neslen) 80° Colton and Tuscher pre-arc subpopulations R=0.99 R=0.90 no data (Figs. 12A, 12B with P = 0.58 from Table 7) R=0.83 indicates that contamination of Colton sand N=7 with reworked Tuscher sand in the Dark Can- R=0.88 R=0.91 n=220 yon sequence could not be readily detected from Castlegate 96° ages of detrital zircons. The greater abundance 97° 105° of arc-derived detrital zircons in the Colton sand N=16 N=7 n=445 n=285 R=0.95 would tend to overprint their lesser proportion in reworked Tuscher sand, which contains a Figure 10. Paleocurrent trends in uppermost Cretaceous and lowermost Paleogene ﬂ uvial higher ratio of subpopulation Ia (Cretaceous) to strata (Fig. 2) along the southern ﬂ ank of the Uinta Basin between the Wasatch Plateau and subpopulation Ib (Jurassic) grains within an arc- the Green River (Fig. 4) keyed to detrital zircon (DZ) chronofacies. Trends depicted are derived subpopulation of the same overall age composite vector means calculated from equally weighted individual vector means for N span (Figs. 9A, 9B). paleocurrent sites (120 total) with 5140 total paleocurrent measurements (n). Rosettes show the radial spread of N vector means but not the full radial spread of n measurements. R is Green River–DeBeque Formations the net resultant vector magnitude with respect to N individual vector means. North Horn in the Wasatch Plateau column (A) includes data from the San Pitch Mountains located farther The age spectra of arc-derived (Fig. 13) and west, closer to the Sevier thrust front (Fig. 3). Colton in the Book Cliffs–Roan Cliffs column pre-arc (Fig. 14) detrital zircon populations in (B) includes data from the Sunnyside delta complex of lowermost Green River Formation stratigraphic units overlying and laterally equiv- (laterally equivalent to uppermost Colton Formation) in Gate Canyon (Figs. 3 and 4) and alent to the Colton Formation delineate the sub- nearby Nine Mile Canyon north of the Roan Cliffs. The Green River column (C) includes regional extent of the Colton arkosic petro facies. data for sub-Colton strata exposed in Range Creek and Tusher Canyon tributary to the Two sandstones from the Green River Formation Green River. The conglomeratic Dark Canyon sequence (Figs. 3 and 4) is present at the (Fig. 3) were collected from interbedded lacus- base of the Colton Formation only east of the Green River (column D). Data are from Chan trine and deltaic strata (Sunnyside delta) capped and Pfaff (1991), Dickinson et al. (1986), Franczyk and Pitman (1987, 1991), Franczyk et al. by the Horse Bench Sandstone Bed (Fig. 4). (1991), Lawton (1983, 1986b), Lawton and Bradford (2011), Lawton et al. (1993), McLaurin They are composited together (sample subset and Steel (2007), Miall and Arush (2001), Olsen (1995), Olsen et al. (1995), Peterson (1976), A of Table 4) as representative of arkosic sand Robinson and Slingerland (1998), Schomacker et al. (2010), Van de Graaf (1972), Willis (Table 2; Fig. 11) that was dispersed into the (2000), and Zawiskie et al. (1982). lacustrine Green River Formation of the Uinta Basin as deltaic distributary channels and mouth

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Qm bars (Schomacker et al., 2010) spread laterally Bluecastle North Horn into Lake Uinta from Colton-like sources. P < “Colton” 0.05 for K-S comparisons of total detrital zircon populations in subset A and the other subsets of Uinta basin DZ samples Castlegate Figure 13 from the Uinta Basin (Table 5), but Sunnyside delta (Gate Canyon) the contrast stems in large part from different Neslen proportions of arc-derived grains, for which Colton (main body) Indianola n~60 P = 0.18–0.49 for the same samples (Table 6). Dark Canyon Arc-derived subpopulations are less abundant sequence DeBeque quartzolithic petrofacies (14%–27%) in Sunnyside delta–Horse Bench (n~295) Tuscher (Fig. 13A) and DeBeque (Fig. 13B) samples Farrer than in the other sample subsets (38%–75%) of Figure 13. P = 0.01 (Table 7) for pre-arc subpopulations (Fig. 14AB) in sample subsets A (Green River) and B (Colton), suggesting that COLTON Dark Green River n=15 Canyon sandstones in the lower Green River Formation n~11 (Sunnyside Tuscher- of the Uinta Basin represent mixtures of Colton- delta in Farrer Price like arkosic detritus (Table 2 and Fig. 11) rich in Gate Canyon) n~45 n=8 River subsurface Green River n~30 arc-derived detrital zircons with pre-arc grains (200–400 m above Colton) Parriette oilfield delivered to Lake Uinta from more proximal F n~30 Lt sources in surrounding Laramide uplifts. For arc-derived grains in sandstones of the Figure 11. Petrofacies relations of Cretaceous and Paleogene sandstones in the Uinta Basin Green River Formation in the Piceance Basin and nearby areas (Fig. 3). See Table 2 for ternary poles (n = number of point counts) and (Davis et al., 2009b), P = 0.00 from K-S com- Figure 2 for stratigraphic relations. Qm—monocrystalline quartz, F—feldspar, Lt—lith- parison with Colton sands of the Uinta Basin but ics. Quartzolithic petrofacies includes Blackhawk, Castlegate, Neslen, Bluecastle, and North P = 0.90 from K-S comparison with DeBeque Horn from the south fl ank of the Uinta Basin, and Currant Creek and Duchesne River from sands of the Piceance Basin (subset C of the north fl ank of the Uinta Basin. Detrital zircon (DZ) samples are from Table 2, Dickinson Table 4), confi rming that sediment transported and Gehrels (2008), and Lawton and Bradford (2011). Other data are from Andersen and into the Uinta Basin by the California paleoriver Picard (1972, 1974), Dickinson et al. (1986), Franczyk and Pitman (1991), Franczyk et al. was not carried across the Douglas Creek arch (1990, 1991), Horton et al. (2004), Lawton (1983, 1986a, 1986b), Lawton and Bradford (2011), into the Piceance Basin (Fig. 1). Note that and Pitman et al. (1982). ca. 50 Ma detrital zircons delivered during

II III IV V 1700 Pre-arc DZ Grains >285 Ma A Colton Formation, Roan Cliffs (Upper Paleocene-Lower Eocene N = 5 n = 195 Figure 12. Comparative age 1450 distribution curves for pre-arc 455 1160 (48%) subset B (older than 285 Ma) detrital zircon (DZ) subpopulations in 1735 Colton (A) and other sample subsets (B–E) from the Utah B Tuscher Formation, Gray Canyon foreland (Table 4) with arc (upper Campanian Cretaceous) index ≥15. Color bands denote N=2 n=131 subpopulations II–V (Table 3). 1085 (72%) Figures on the right denote the subset D 1705 percentages of pre-arc grains C Farrer Formation, Gray Canyon 1075 in the total grain population (upper Campanian Cretaceous) of each sample subset. N is the N=2 n=97 (52%) number of samples composited subset E 1085 for each age spectrum and n is D Kaiparowits Formation 1710 the number of dated grains in Kaiparowits Plateau (26%) (upper Campanian) subset F each composite DZ population. N = 5 n = 103 1075-1160 1695 E pre-upper Campanian Utah foreland 1450 (85%) N = 6 n = 443 subsets K and L

250 500 750 1000 1250 1500 1750 2000 2250 Age (Ma)

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Ia Ib Ic Tuscher (Fig. 13E) and Flagstaff (Fig, 13D) ARC DZ GRAINS 86 170 <285 Ma 80 165 samples are absent from the Colton samples A. Green River Formation (Fig. 13C), but are also prominent for the Kai- (Lower Eocene) parowits Formation of the Table Cliff basin (Fig. Gate Canyon, Uinta Basin 9D). The similarity of Kaiparowits and Farrer– (14%) N=2 n=26 subset A Tuscher detrital zircon populations can be attrib- 77 uted to an upstream-downstream relationship between Kaiparowits and Farrer depositional systems (Lawton and Bradford, 2011). Field evidence for possible reworking of Kaiparowits 66 B DeBeque Formation sand into the Flagstaff depositional system is 156 (Paleocene-Eocene) 174 Piceance Basin hidden beneath the Marysvale volcanic fi eld (27%) N = 2 n = 33 (Fig. 1), but feldspathic sandstones (Stanley and subset C 168 Collinson, 1979) in the Flagstaff Limestone of the Wasatch Plateau north of the volcanic fi eld C Colton Formation may refl ect a sedimentological connection. 94 (Paleocene-Eocene) Uinta Basin (52%) N = 5 n = 221 subset B COLTON RECYCLING OPTION 164 172 D Flagstaff Limestone Based on his interpretation that the California 91 (Paleocene-Eocene) 72– paleoriver reversed course through the Grand 74 Soldier Summit N = 1 n = 53 Canyon region (Fig. 1) from northeastward (75%) Cretaceous paleofl ow to southwestward Paleo- GRF gene paleofl ow, Wernicke (2011) suggested that 152 Colton detritus was recycled from Kaiparowits- E Farrer and Tuscher equivalent strata exposed in Laramide uplands 76 98 Formations (upper Campanian) of southern Utah; his hypothesis cannot be (38%) Gray Canyon tested directly because the only Kaiparowits- subsets N = 4 n = 144 D and E equivalent strata known south of the Uinta Basin in the Utah foreland, other than exposures of the 0 50 100 150 200 250 300 Age (Ma) Kaiparowits Formation in the Table Cliff basin (Fig. 15), form a thin succession (~30 m) of fi ne- Figure 13. Comparative age distribution curves for arc-derived (younger than 285 Ma) grained strata at the erosional top of the Creta- detrital zircon (DZ) subpopulations in Colton (C) and related samples from the Uinta Basin ceous section preserved in the Henry Mountains and nearby areas with arc index ≥12. Color bands denote subpopulations Ia, Ib, and Ic basin (Fig. 1). Detrital zircon populations, sand- (Table 3). Figures on the left denote the percentages of arc-derived grains in the total grain stone petrofacies, paleocurrent trends over time, population of each sample or sample subset. N is the number of samples composited for each relative sediment volumes, and the timing of age spectrum, and n is the number of dated grains in each DZ population plotted. Kaiparowits stripping off the Kaibab uplift do not favor recycling of Colton arkosic sediment from Cretaceous strata. the interval 49.5–47.0 Ma to the Green River that are absent from Colton (Fig. 13C) grains Basin north of the Uinta uplift (Figs. 1 and 3) probably refl ect derivation of Cretaceous detrital Detrital Zircons by the Idaho paleoriver (Chetel et al., 2011) zircons in DeBeque sand from igneous rocks of from sources in the Challis volcanics were not the Colorado Mineral Belt (Fig. 1), where ages Colton and Kaiparowits detrital zircon popu- detected in the arkosic Horse Bench Sandstone of 75–65 Ma for intrusions are characteristic lations yield P = 0.00 for total, arc-derived, Bed (ca. 47.5 Ma) at Gate Canyon (Table 2; (Chapin, 2012). The Yavapai-Mazatzal age peak and pre-arc grains (Tables 5–7). This result is Fig. 4), or in sandstones from the Green River for DeBeque (subpopulation V of Fig. 14B) was atypical for K-S analysis of recycled detrital Formation in the Piceance Basin (Davis et al., probably derived from Precambrian basement in zircons and their source strata. For example, 2009b). Volcaniclastic detritus derived from the Laramide uplifts of Colorado rather than hav- P = 0.59 for pre-arc subpopulations in Lower north (Surdam and Stanley, 1980) evidently did ing Mogollon provenance, and the DeBeque and Middle Jurassic eolianites of the western not reach basins south of the Uinta uplift until Yavapai-Mazatzal age peak is slightly older Colorado Plateau and in Upper Jurassic sand- after ca. 47.5 Ma. than the Yavapai-Mazatzal age peaks for Colton stones of the Morrison Formation derived from P ≤0.01 for K-S comparisons of total, arc- and related strata in Utah. The DeBeque pre-arc recycling of the older eolianites from the Sevier derived, and pre-arc detrital zircons in Colton age subpeak at 1885 Ma is not present on any of thrust belt (Dickinson and Gehrels, 2008), and and DeBeque sample subsets (Tables 5–7), in the age spectra for strata deposited in Utah (Fig. P = 0.69 for pre-arc subpopulations in Middle keeping with fl uvial paleofl ow to the northwest 14), and may be a signal of Colorado rather to Upper Jurassic eolianites of the eastern out of Colorado for the DeBeque Formation than Mogollon provenance within the Yavapai- Colorado Plateau and in Lower Cretaceous (Lorenz and Nadon, 2002), rather than from Mazatzal belt. sandstone of the Cintura Formation in south- Utah to the west. The prominence of 60–80 Ma Subpeaks in the range of 70–85 Ma for arc- eastern Arizona derived from recycling of older age subpeaks for DeBeque grains (Fig. 13B) derived subpopulation 1a (Table 3) in Farrer– eolianites uplifted along the Mogollon high-

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II III IV V

A Green River Formation (Lower Eocene) 1720 Pre-arc DZ Grains >285 Ma Gate Canyon, Uinta Basin n = 2 n = 154 1080 (86%) subset A Figure 14. Comparative age 1740 distribution curves for pre-arc (older than 285 Ma) detrital B DeBeque Formation, Piceance Basin zircon (DZ) subpopulations in 1885 (Paleocene-Eocene) N = 2 n = 90 (73%) Colton (C) and related samples 1495 subset C from the Uinta Basin and envi- 1700 rons with arc index ≥12. Color C Colton Formation, Uinta Basin bands denote subpopulations (Paleocene- Eocene) n = 5 n = 195 1450 (48%) II–V (Table 3). Figures on the subset B right denote the percentages of 1695 pre-arc grains in the total grain population of each sample or sample subset. N is the num- D Flagstaff Limestone, Soldier Summit ber of samples composited for (Paleocene-Eocene) n = 1 n = 18 each age spectrum, and n is the 1435 (25%) number of dated grains in each GRF DZ population plotted. Pre- 1695 arc grains older than 2250 Ma E Farrer and Tuscher Formations, Gray Canyon (3.5% of the total) were omit- 1410 (upper Campanian Cretaceous) (62%) ted for reasons of scale. N=4 n=228 subsets D and E

250 500 750 1000 1250 1500 1750 2000 2250 Age (Ma)

lands rift shoulder of the Bisbee basin (Dickin- ous rocks, and that the higher proportion of the umes ~3000 km3 for the Colton versus ~500 km3 son et al., 2009). Permian–Triassic subpopulation Ic in Colton for the Kaiparowits. All preserved Kaiparowits Arc-derived grains are more abundant in Kai- sand refl ects derivation from farther south and Formation remnants were buried beneath latest parowits samples (50%–90%) than in Colton east, where Permian and Triassic components Cretaceous and earliest Paleogene strata at the samples (44%–58%), with a net Kaiparowits of the Cordilleran igneous assemblage are more time of Late Paleocene–Early Eocene Colton content of 74% and a net Colton content of 52% prominent. sedimentation (Fig. 15). The hypothesis that (Fig. 9). Mogollon indices are 66 for Colton but Kaiparowits sand was recycled into the Colton only 41 for Kaiparowits (Figs. 6 and 12), refl ect- Sandstone Petrofacies Formation thus requires the supposition that ing derivation of Kaiparowits pre-arc subpopu- Kaiparowits-equivalent strata were once pres- lations dominantly from the Sevier thrust belt to Kaiparowits sandstones are uniformly more ent over large areas to the east of the Table the west with minimal input from the Mogollon lithic than Colton sandstones (Fig. 17). The Cliff syncline, where they were not capped by highlands tapped by the California paleoriver largely volcaniclastic Kaiparowits petrofacies younger strata but were later removed by Paleo- farther south (Fig. 1). Paleocurrent trends in the contrasts more strongly with the Colton arkosic gene erosion. Kaiparowits Formation are eastward (Fig. 16), petrofacies than does any other Cretaceous– Kaiparowits strata in the Table Cliff syncline refl ecting derivation of detritus from near the Paleogene sandstone suite of the Utah fore- constitute a distal fl uvial facies of meander belt southern terminus of the Sevier thrust belt rather land other than the quartzolithic petrofacies and anastomosed alluvial deposits (Eaton et al., than from the central Mojave headwaters region derived from recycling of sedimentary detritus 1987; Goldstrand, 1992; Lawton et al., 2003; for the California paleoriver (Fig. 1). from the Sevier thrust belt (Figs. 11 and 17). Roberts, 2007; Lawton and Bradford, 2011). For arc-derived grains, the principal subpeaks Destruction of 50%–75% of lithic fragments in Only the lowermost Kaiparowits Formation for subpopulation Ia (Cretaceous) are 75–80 Ma sand by prefer ential abrasion during recycling, preserves any record of tidal infl uence on sedi- for Kaiparowits grains but ca. 95 Ma for Colton while maintaining the quartz to feldspar ratio mentation from the Cretaceous interior seaway, grains, and for subpopulation Ib (Jurassic) are unchanged, seems unlikely and an unnecessary and the bulk of the unit may have formed as a 152 Ma for Kaiparowits grains and 168 Ma for postulate given the comparative Kaiparowits- terminal fl uvial accumulation deposited rapidly Colton grains (Figs. 9A, 9D). The content of Colton detrital zircon data. within an incipient Laramide down-bowing on subpopulation Ic (Permian–Jurassic) is propor- the site of the Table Cliff syncline. Kaiparowits- tionally four times as great in Colton sand (16% Sediment Volumes equivalent strata in the Henry Mountains basin of subpopulation I) as in Kaiparowits sand (4% to the east are gray mudstone with minor thin of subpopulation I). We infer that arc-derived The Kaiparowits Formation in the Table Cliff sandstone interbeds (called “beds on Tarantula Cretaceous and Jurassic detrital zircons in Kai- syncline is approximately as thick (~1000 m) as Mesa” by Peterson and Ryder, 1975) that over- parowits and Colton sands had their sources in the Colton Formation at Desolation Canyon of lie the Tarantula Mesa Sandstone (Eaton, 1990; different segments of the Cordilleran magmatic the Green River (Fig. 4), but is less extensive Roberts et al., 2005; Jinnah and Roberts, 2011). arc exposing a different mix of Mesozoic igne- laterally (Fig. 1), with preserved sediment vol- The Tarantula Mesa Sandstone is a correlative

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Figure 15. Uppermost Cre- B-R taceous (Campanian–Maas- Sevier CIRCLE thrust trichtian) and lower Paleogene front Pa CL CLIFFS (Paleocene–Eocene) strata w Pu K Tc UPLIFT (Fig. 2) in the Table Cliff syn- Cenozoic CB Tc E cline (TCS) of the Kai paro- zone of volcanic CP-PH wits Plateau and on adjacent complex pCK pKs faulting CA high plateaus (Paunsagunt- CC cover B-R PTS pCUK PH Markagunt) westward to the K5 edge of the Basin and Range T t l W3 Province (B-R, hachured line) u K1 a H adapted after Hintze (1963), f Tc 37° pCUK Tc t T Hackman and Wyant (1973), l 30′ er i u F W1 N I Doelling and Davis (1989), v a pKs e f L e TCS Kk Goldstrand (1990, 1992, 1994), S P n i Kw pCUK l U c Bowers (1991), Taylor (1993), Kw o n pKs o Goldstrand and Mullett (1997), m G t B

n A Kw Eaton et al. (2001), Lawton b u pKs B a g I b N i et al. (2003), Sable and Here- a a A K PAUNSAGUNT s K ford (2004), and Larsen et al. MARKAGUNT n t u PLATEAU s (2010). Sevier thrust front a pCUK PLATEAU E (largely masked by cover) after Pa pKs Anderson and Dinter (2010). selected towns 25 km Paunsagunt thrust system 113° W DZ samples 112° W KAIPAROWITS PLATEAU (PTS), induced by the gravi- metric load of the Oligocene Tc CP-PH Kk Kw pCUK pKs Marysvale volcanic ﬁ eld to the Claron Canaan Peak and Kaiparowits Wahweap pre-Campanian pre- Formation Pine Hollow Formation Formation Upper Cretaceous northwest, includes the master (Eocene) Formations (Campanian) (Campanian) Cretaceous strata south-vergent Rubys Inn thrust and the linked Pine Hill backthrust (Lundin, 1989; Merle et al., 1993). Grand Castle Formation is interpreted as a proximal facies of the capping sandstone member of the Wahweap Formation (Lawton et al., 2003; Hunt et al., 2011; Johnson et al., 2011). Strata above the Kai parowits Formation in the Table Cliff syncline depicted locally as Grand Castle Formation (Larsen et al., 2010) are here regarded as Canaan Peak Formation (Bowers, 1972). The outcrop width of the Wahweap Formation is exaggerated locally for clarity on the Paun- sagunt and Markagunt Plateaus where its exposed thickness is ≤150 m (Bowers, 1991; Lawton et al., 2003; Jinnah and Roberts, 2011). Detrital zircon (DZ) samples from the Table Cliffs syncline (see Table 1): CA (sample 5JL5) and CB (sample CP1)—Canaan Peak (n = 2), CL—Claron (n = 1), K1—lower Kaiparowits (n = 1), K5—middle and upper Kaiparowits (n = 5), PH—Pine Hollow (n = 4), W1—lower member of Wahweap Formation (n = 1), W3—upper and capping sandstone members of Wahweap Formation (n = 3). Towns: CC—Cedar City, E—Escalante, G—Glendale, H—Henrieville, Pa—Parowan, Pu—Panguitch, T—Tropic.

of the capping sandstone member of the Wah- mal facies of the Kaiparowits Formation once the capping sandstone member of the Wahweap weap Formation (Figs. 2 and 17) that under- extended westward across the Laramide Kaibab Formation underlying the Kaiparowits Forma- lies the Kaiparowits Formation in the Table uplift toward the Sevier thrust front ~100 km tion are preserved as the conglomeratic Grand Cliff syncline. No sandstone-rich Kaiparowits from the Table Cliff syncline (Fig. 15). A poten- Castle Formation (Lawton et al., 2003; Johnson equivalents are known from any locale east of tial means for recycling Kaiparowits sand into et al., 2011), which was long regarded as Paleo- the Table Cliff syncline and south of the Uinta younger units during Laramide deformation gene in age (Goldstrand, 1990, 1992, 1994; Basin, where the Farrer Formation contains is to posit removal of Kaiparowits strata from Goldstrand et al. 1993; Goldstrand and Mullett, similar detrital zircon subpopulations (P = 0.55 the crest of the Kaibab uplift, which displays 1997; Goldstrand and Eaton, 2001) before the for arc-derived grains and P = 0.36 for pre-arc 1.6 km of structural relief adjacent to the Table recent discovery of Campanian palynomorphs grains; Tables 6–8). Cliff syncline (Tindall et al., 2010). and a dinosaur track in its middle sandstone Thin remnants of Kaiparowits Formation are member (Hunt et al., 2011). The oldest strata Kaibab Erosion exposed locally on the Paunsagunt Plateau near lapping across the northern end of the eroded the crest of the Kaibab uplift (Fig. 15), where Kaibab uplift toward the relict Sevier thrust The distal character of Kaiparowits fl uvial they conformably overlie middle Cam panian front are lacustrine and associated fl uvial strata facies in the Table Cliff syncline, paleocurrents Wahweap Formation and unconformably under- of the Eocene (and possibly Upper Paleocene) from the southwest, and detrital zircons refl ect- lie the Paleogene Claron Formation (Bowers, Claron Formation (Fig. 2), but stratigraphic ing derivation of Kaiparowits sand from both 1991), but are not known on the Markagunt relations within the Table Cliff syncline pre- the Sevier thrust belt and elements of the Cordi- Plateau farther west (Eaton et al., 2001; Moore clude transport of Kaiparowits detritus toward lleran magmatic arc beyond or near the southern and Straub, 2001; Lawton et al., 2003; Roberts, the Uinta Basin from the Kaibab uplift during terminus of Sevier thrusting imply that proxi- 2007). On both high plateaus, proximal facies of Paleocene–Eocene Colton sedimentation.

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A B C peak Laramide deformation (Fig. 2). Subrounded Stratal Markagunt Paunsagunt Table Cliff units conglomerate cobbles are dominantly mature Plateau Plateau syncline quartzite, chert-argillite, and igneous felsite N = 5 R = 0.76 R = 0.67 N = 3 (Bowers, 1972; Schmitt et al., 1991; Goldstrand, n = 185 n = 100 1992). The lack of any sources for those clast 86° 88° Claron types in Kaiparowits or subjacent Cretaceous units of the Kaibab uplift implies delivery of N = 5 the clasts directly to the Canaan Peak braidplain 146° n = 165 R = 0.40 KEY from the Sevier thrust front or beyond (Schmitt N = 12 et al., 1991). This observation, coupled with n = 290 percentage of angularity of ~10° at the Kaiparowits–Canaan Pine arc-derived Hollow non-deposition non-deposition DZ grains Peak unconformity (Bowers, 1972), suggests (<285 Ma) that Kaiparowits strata had already been stripped from the Kaibab uplift by Maastrichtian or Early 196° R = 0.32 50%–90% Paleocene time. A sandstone sample from the N = 5 56° Canaan Peak Formation (Table 1) contains n = 220 <10% arc-derived detrital zircon grains (Larsen Canaan non-deposition et al., 2010), showing that Canaan Peak sand is Peak 15%–30% 89° N = 17 not reworked Kaiparowits sand, which contains n = 900 R = 0.81 50%–90% arc-derived detrital zircon grains. R = 0.64 Overlying the Canaan Peak Formation con- N = 17 5%–15% n = 310 70° cordantly along the axis of the Table Cliff syn- non-deposition Laramide cline (Bowers, 1972; Goldstrand, 1990), but o r erosion Kaiparowits Laramide (l o c a l with angularity of 5°–10° (Goldstrand, 1990; erosion remnants) <5% Larsen et al., 2010) in the limbs of the syncline, R = 0.79 the Pine Hollow Formation (80–120 m thick) is partly equivalent in age to the Colton For- R = 0.97 Grand no data mation (Fig. 2) but displays centripetal paleo- Castle R = 0.84 ( = ) currents inward toward the Table Cliff syncline Wahweap with a weak resultant vector southward along capping N = 6 N = 7 sandstone n = 450 the syncline axis (Fig. 16). The paleocurrrent 123° n = 245 member 130° implication of closed interior drainage for the N = 6 Pine Hollow Formation within the deforming 5° R = 0.95 Wahweap n = 285 31° Table Cliff syncline, bounded by active Kaibab (upper, and Circle Cliffs Laramide uplifts (Goldstrand, middle, lower N = 3 1992, 1994; Larsen et al., 2010), is confi rmed members) n = 65 by Pine Hollow facies patterns. Characteristic R = 0.97 cyclic sedimentation encompassed four to six cycles, each fi ning and drying upward from Figure 16. Paleocurrent trends in uppermost Cretaceous and lower Paleogene fl uvial strata alluvial fan to playa lake deposits (Larsen et al., (Fig. 2) in the Table Cliff syncline and plateaus (Paunsagunt, Markagunt) to the west (see 2010) in successions typical of undrained sedi- Fig. 15) keyed to detrital zircon (DZ) chronofacies. Trends depicted are composite vector mentary basins. Early Paleocene palynomorphs means calculated from equally weighted individual vector means for N paleocurrent sites in the basal Pine Hollow Formation (Goldstrand, (85 total) with 3215 total paleocurrent measurements (n). Rosettes show the radial spread 1990, 1994; Goldstrand and Eaton, 2001), and of N vector means but not the full radial spread of n measurements. R is the net resul- Late Paleocene to Early Eocene palynomorphs tant vector magnitude with respect to N individual vector means. Note that strata referred at higher horizons (Larsen et al., 2010), indicate by Goldstrand (1990) to Canaan Peak Formation on the southeastern Paunsagunt Plateau that no Kaiparowits or other detritus could have were mapped as part of the Claron Formation by Bowers (1991), and that net Canaan Peak exited through the Table Cliff synclinal basin and Claron paleocurrent vectors are nearly identical for the Paunsagunt Plateau. Data are during the time frame of Colton sedimentation from Goldstrand (1990), Goldstrand and Mullett (1997), Jinnah et al. (2009), Larsen et al. in the Uinta Basin (Fig. 2). (2010), Lawton et al. (2003), Roberts (2007), and Schmitt et al. (1991). Table Cliff Strata

In the Table Cliff syncline, the Kaiparowits 16). Campanian palynomorphs in the Canaan The last phase of Paleogene sedimentation Formation is unconformably overlain by the con- Peak Formation are probably reworked (Eaton, recorded by strata in the Table Cliff syncline glomeratic Canaan Peak Formation (80–140 m 1991; Goldstrand, 1992, 1994; Goldstrand et al., and across the Kaibab uplift was deposition of thick), which was deposited on a coarse alluvial 1993) because abundant Early Paleocene palyno- lacustrine and associated ﬂ uvial strata in the braidplain (Bowers, 1972; Schmitt et al., 1991; morphs (Goldstrand, 1990) suggest syntectonic Claron Formation, which formed an Eocene Larsen et al., 2010) by eastward paleoﬂ ow sub- Canaan Peak sedimentation during the post-Kai- stratigraphic cap (Eaton et al., 2011) over both parallel to Kaiparowits paleocurrent trends (Fig. parowits Maastricthian–Paleocene time frame of structural features when Laramide deformation

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Qm neared completion (Fig. 15). Local incision of Straight Cliffs capping Wahweap paleosols in marginal-lacustrine fl uvial deposits n~60 record fl uctuations in Claron lake level (Gold- Dakota strand and Eaton, 2001) that may have refl ected Utah foreland DZ samples Buckhorn waning phases of Laramide deformation. Kaiparowits upper Wahweap Detrital zircon populations of post-Kaiparo- Wahweap Ferron Ferron wits strata in the Table Cliff syncline (Fig. 18A) refl ect minor addition of arc-derived grains (arc Ferron quartzolithic Dakota petrofacies indices 7–17) to pre-arc grains either derived Buckhorn (n~125) from the Sevier thrust belt (Mogollon indi- COLTON ces 24–36) or recycled from Cretaceous strata n=15 derived ultimately from the Sevier thrust belt (Larsen et al., 2010). Their age spectra resemble lower and middle middle Campanian (pre-Kaiparowits) Upper Wahweap n~10 Cretaceous strata of the Kaiparowits Plateau and high plateaus to the west (Fig. 18B), and are not markedly different from pre-middle Cam- middle and upper panian strata exposed farther east in the Henry Kaiparowits n~15 Mountains basin (Fig. 1) and nearby areas (Fig. 18C). All have Mogollon indices <45 (Table 3; F Lt Fig. 7). The Laramide Table Cliff basin was evidently screened from the California paleoriver Figure 17. Petrofacies relations of Cretaceous and Paleogene sedimentary strata in the Table by the Circle Cliffs uplift throughout Maastrich- Cliff syncline and nearby areas (Fig. 15). Qm—monocrystalline quartz, F—feldspar, Lt—lithics . tian–Paleogene time (Fig. 1), and lacked any See Table 2 for defi nitions of ternary poles (n = number of point counts) and Figure 2 for strati- sedimentary connections to wider reaches of the graphic relations. Colton fi eld (for comparison) is from Figure 11. Quartzolithic petrofacies foreland region. includes upper and capping Wahweap and basal Kaiparowits from the Kaiparowits Plateau, upper and capping Wahweap from the Paunsagunt and Markagunt Plateaus, Canaan Peak LARAMIDE SEDIMENT TRANSPORT and Pine Hollow from the Table Cliff syncline, and Claron from across its outcrop belt on the High Plateaus. Detrital zircon (DZ) samples are from Table 2, Dickinson and Gehrels (2008), Patterns of detrital zircon ages, paleocurrents, and Lawton and Bradford (2011). Other data are from Allen and Johnson (2010), Goldstrand and petrofacies in strata of the Utah foreland (1992), Jinnah et al. (2009), Larsen et al. (2010), Lawton et al. (2003), and Schmitt et al. (1991). suggest that intraforeland uplifts formed during

I II III 1V V 100

80 A subset G and 12JL5 170 410 1050 Canaan Peak - Pine Hollow - Claron 600 1425– (Maastrichtian - Paleocene - Eocene) 1455 of Table Cliff basin N=6 n=280 Figure 18. Comparative age distribution curves for detrital zir- 415 B 80 1095 subset J con (DZ) populations in Maas- middle Campanian of Markagunt - 595 trichtian–Paleogene strata of 155 Paunsagunt - Kaiparowits Plateaus the Table Cliff basin (A) and (southern Sevier foredeep deposits) N = 5 n = 403 older Cretaceous strata of the Sevier foredeep (B, C). Color 96 bands denote subpopulations I–V (Table 3). N is the number of samples composited for each age spectrum, and n is the num- 160 ber of dated grains in each composite DZ population. C subsets K and L pre-middle Campanian Upper Cretaceous of Utah foreland basin 1085 1700 N=6 n=570 425 1450

0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 3250 Age (Ma)

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Laramide deformation forced changes in the con- IIVV fi guration of fl uvial systems carrying sediment 160– 1705 A 175 Chuska Sandstone longitudinally northward east of the Sevier thrust 3% I 1425 (Oligocene) front. In the Uinta Basin, the pebbly beds present N = 1 n = 100 locally as a lateral equivalent of the Tuscher For- 1695 mation and the Dark Canyon sequence above the B Tuscher Formation provide a diachronous record Music Mountain Formation of incipient Laramide deformation, progress- 1405 (Eocene) ing from west to east and in time producing the 100 5% I N=2 n=196 Uinta Basin, where nearly 5000 m of Paleogene 170 sediment accumulated (Johnson and Johnson, 1991). Before mid-late Campanian time (Fig. 2), 168 data are consistent with master streams fl ow- C ing along the axis of the Sevier foredeep within Colton Formation 94 52% I ~100 km of the thrust front and directly over the (Paleocene - Eocene) N = 5 n = 426 sites of the younger Kaibab, Circle Cliffs, and 225 1705 San Rafael uplifts (Fig. 1). Beginning in lat- 1460 est Campanian time and continuing into early 166 Paleogene time, the growth of Laramide uplifts blocked sediment transport and shifted longitudinal sediment dispersal through the foreland region to a pathway east of the uplifts. Growth faults along the East Kaibab mono- 96 51% I cline fl anking the Kaibab uplift bracket the onset D of Laramide deformation at 80–76 Ma (Tindall Dome Rock succession (Late Cretaceous) et al., 2010) in middle Campanian time and thin- N=1 n=99 ning of Cretaceous strata across the San Rafael Swell farther north document initiation of that 1450 1745 uplift as a subsurface growth fold as early as 235 77 Ma (Aschoff and Steel, 2011) within the same 0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 time frame. Initiation of the uplifts as buried sub- Age (Ma) surface structures did not initially preclude transport of sediment across their evolving crests, Figure 19. Comparative age distribution curves for detrital zircon but the unconformity between the Kaiparowits (DZ) populations in Cretaceous-Paleogene strata of northern Ari- Formation and the Canaan Peak Formation in zona (A, B, D) and the Colton Formation of the Uinta Basin (C). the Table Cliff syncline implies that the Kaibab Color bands denote subpopulations I and IV–V (Table 3). N is the uplift had attained surfi cial relief by Maastrich- number of samples composited for each age spectrum, and n is tian time. Ponding of the Maastrichtian to Lower the number of dated grains in each DZ population plotted. North- Paleocene North Horn Formation within a depo- ern Arizona samples (see Table 1): A—NARCHU, B—COL12 and center west of the San Rafael Swell (Lawton, COL13 (composited), D—DR. 1983, 1986b) implies the existence of sediment- blocking surfi cial relief on that uplift during the same time frame. We infer that diversion of fl u- but others are dissimilar, raising unresolved arc-derived grains, 2 are Cretaceous (90–99 Ma) vial paleofl ow in mid-late Campanian time from questions about the confi guration of Laramide and 8 are Jurassic (158–173 Ma), but could a Kaiparowits–Farrer pathway across the sites paleodrainage patterns south of the Utah fore- derive from Mesozoic plutons intrusive into of Laramide uplifts to the younger California land. In an unsuccessful attempt to intercept Yavapai-Mazatzal basement of Arizona rather paleoriver east of the uplifts for later Tuscher– Mojave-derived sediment in transit toward the than from arc assemblages in California. Detri- Colton sedimentation refl ected the tectonically Uinta Basin along an upstream reach of the tal zircon populations from the Music Mountain induced adjustment of the confi guration of the California paleoriver, we collected two samples Formation resemble those from Oligocene eoli- master fl uvial systems in the Utah foreland to (COL12 and COL13; see Table 1) from the anite of the Chuska Sandstone (Figs. 1 and 19A) Laramide deformation. Laramide drainages were Paleogene Music Mountain Formation (Young, derived from the defl ation of alluvial deposits forced to thread pathways between Laramide 1999) south of the Grand Canyon in northwest- spread northward from the Mogollon region of foreland uplifts that grew well to the east of the ern Arizona (Fig. 1). The 2 samples contain central Arizona (Dickinson et al., 2010). Recy- Sevier thrust front (Fig. 1). only 5% arc-derived detrital zircon grains and cling of Music Mountain sand into Chuska sand none of Paleozoic–Neoproterozoic or Grenville is not favored, however, by K-S analysis yield- ARIZONA RELATIONS age (subpopulations II and III of Table 3), with ing P = 0.01 for comparison of Music Mountain 99.5% of the pre-arc detrital zircons derived and Chuska detrital zircon populations. Some detrital zircon populations from Cre- from Yavapai-Mazatzal and anorogenic granite Failure to detect a Colton-like detrital zircon taceous–Paleogene strata in northern Arizona sources (subpopulations IV and V of Table 3) of population in the Music Mountain Formation are similar to the Colton population (Fig. 19), the Mogollon highlands (Fig. 19B). Of the 10 (Fig. 19B) is explicable in any one of four ways,

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the fi rst two spatial and the latter two temporal. no arc-derived detrital zircons reached the Uinta drainages (Fig. 1). By contrast, backarc succes- (1) Music Mountain paleodrainages were con- Basin during North Horn sedimentation (Fig. 4). sions of Utah received signifi cant components tiguous northward with paleodrainages on If the Music Mountain Formation is slightly of Jurassic grains (older than 150 Ma) from east the west side of the Kaibab uplift and did not younger (Middle or Late Eocene), as inferred of the paleodrainage divide throughout Late connect with the California paleoriver fl owing by Cather et al. (2008), its deposition postdated Cretaceous time (Fig. 20E, 20G), although com- east of the Kaibab uplift. (2) Music Mountain the delivery of voluminous arc-derived detrital parative data are unavailable for Maastrichtian paleodrainages were local tributaries to the main zircons to the Uinta Basin, and the paucity of backarc successions. stem of the California paleoriver and did not arc-derived detrital zircons in the Music Moun- For Paleogene successions (Figs. 20A, 20B), tap Mojave sources. (3) Aggradation of Music tain Formation presents no conceptual problem. Cretaceous and Jurassic subpopulations of Mountain paleodrainages occurred during arc-derived grains are prominent in both back- Laramide deformation when delivery of sedi- Dome Rock Succession arc (Colton) and forearc strata, suggesting that ment from the Mojave region to the Uinta Basin the Paleogene drainage divide had migrated was temporarily interrupted during adjustment The age spectrum of detrital zircons in the far enough inland by Middle Paleocene time of drainage patterns to the growth of Laramide Upper Cretaceous (younger than 80 Ma) Dome for Cretaceous and Jurassic elements of the uplifts. (4) Music Mountain strata are slightly Rock succession (Spencer et al., 2011) of the magmatic arc assemblage to contribute detri- younger than the Colton Formation and accu- upper McCoy Mountains Formation (Fig. 19D), tus to both forearc and backarc paleodrainages mulated after the delivery of Mojave detritus to exposed near the Colorado River in western Ari- (Fig. 1). Minor differences between the forearc the Uinta Basin. zona (Fig. 1), yields P = 0.35 from K-S compari- and backarc assemblages are diffi cult to evalu- The Music Mountain Formation fi lled paleo- son with the Colton age spectrum (Fig. 19C). ate, but it is notable that backarc Colton (Fig. valleys trending generally from south to north, Arc-derived grains derived from the Mojave 20B) incorporates more of subpopulation Ic and also prograded northward as associated segment of the Cordilleran magmatic arc form (Permian–Triassic ) derived mainly from Mexi- alluvial aprons spread across northwestern essentially the same proportion of detrital zir- can segments of the magmatic arc (Table 3), and Arizona south of the modern Grand Canyon con grains in each case, with the pre-arc grains that forearc Eocene strata (Fig. 20A) include an (Young, 1999, 2001a, 2001b). Continuation derived predominantly from subpopulations Early Cretaceous subpeak at 117 Ma not present of the Music Mountain paleodrainage system IV and V (Table 3) of Mogollon provenance. for the backarc. The general similarity of forearc northward across the site of the modern Grand Minor detrital zircon grains in the age range and backarc Paleogene curves is consistent, Canyon (Young, 1982, 1985; Graf et al., 1987) of 250–1250 Ma in the Colton Formation, but however, with derivation of arc-derived detritus would have carried its detritus to the western absent from the Dome Rock succession, are in the Colton Formation largely from the Mojave fl ank of the Kaibab uplift in Utah (Fig. 15). inferred to refl ect admixture of subordinate segment of the Cordilleran magmatic arc. From that position, Music Mountain sediment detritus from the Sevier thrust belt into the Detrital zircon populations derived during could not have crossed the enclosed Paleocene– California paleoriver during sediment transit Paleogene time from different longitudinal seg- Eocene (Fig. 2) Pine Hollow depositional system northward through the Utah foreland. Dome ments of the Cordilleran magmatic arc are simi- in the Table Cliff syncline to join the Califor- Rock and Colton detrital zircon populations link lar from the Mojave region northward along the nia paleoriver farther east, and can be treated as both sedi mentary successions to sources in the Sierra Nevada (Fig. 21). The age spans of the the deposits of a local fl uvial system separate Mojave region, but the two successions are not principal subpopulations document similar ages from the California paleoriver system. If the coeval and provide no information on the evolu- of igneous source rocks along the full length of Music Mountain paleodrainage instead curved tion of drainage patterns near the Mojave region the Sierra Nevada–Peninsular Ranges segment eastward as a local tributary of the California during intervening Laramide deformation. of the Cordilleran magmatic arc. The north- paleoriver system fl owing past the southern end ern, central, and southern Sierra Nevada curves of the Kaibab uplift (Fig. 1), its sediment would FOREARC-BACKARC COMPARISON were composited for auriferous gravel samples have enhanced the volume of Mogollon detritus from the Yuba paleodrainage (Fig. 21A) and in the pre-arc subpopulations of Colton detrital The detrital zircon data of Jacobson et al. the Mokelumne-Stanislaus paleodrainages zircons (Fig. 12A) without contributing non- (2011) for Cretaceous–Paleogene sedimentary (Fig. 21B), and from sedimentary assemblages arc sediment voluminous enough to dilute the arc assemblages exposed in the southern California exposed in the El Paso and San Emigdio Moun- signature of Mojave-derived detrital zircons in forearc allow comparison with data from the tains (Fig. 21C) aligned along the southern or Colton sand. Utah foreland backarc to assess the character Tehachapi tail of the Sierra Nevada block. The Music Mountain Formation is judged of arc-derived sediment contributed from the All the age spectra of Figure 21 display Cre- from its nonmarine molluscan fauna to be of Mojave region to opposite sides of a coastal- taceous and Jurassic peaks (subpopulations Ia Early Eocene age (Young and Hartman, 2011; inland drainage divide within the Cordilleran and Ib of Table 3) separated by troughs repre- Young et al., 2011), coeval with the Colton magmatic arc (Fig. 1). The forearc successions senting the regional null in arc magmatism. The Formation, meaning that one of the two spa- used for the assessment are those deposited in Sierra Nevada curves (Figs. 21A–21C) display tial explanations for the dichotomy between forearc basins of the western Salinian block, the age subpeaks of ca. 115 Ma and 145–150 Ma Music Mountain and Colton detrital zircons is Sur-Nacimiento block, and the western Trans- that are not present for Colton (Fig. 21D) or for the favored rationale. Other ages for the Music verse Ranges (Fig. 1). southern California forearc assemblages (Figs. Mountain Formation are considered here only Pre–Middle Paleocene forearc successions 21E, 21F). That observation supports the infer- because it has to date yielded no defi nitive mam- (Figs. 20C, 20D, 20F) are dominated by Creta- ence that Colton sand included important con- malian fauna. If the Music Mountain Formation ceous (younger than 120 Ma) grains, suggesting tributions of arc-derived grains from the Mojave proves to be slightly older (Paleocene in age), it that the drainage divide was then close to the region far to the south, but not from the Sierra may have been deposited during the Laramide coast, with Jurassic arc assemblages exposed Nevada directly to the west. The presence of interval of foreland drainage readjustment when farther east beyond the reach of coastal paleo- an age subpeak at 117 Ma for Eocene forearc

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ARC-DERIVED 101 A. Lower to Middle Eocene 97 [arc-derived DZ subpopulations (50–282 Ma)] DZ GRAINS 76 117 162 forearc strata (<285 MA) 170 N=19 n=298

168 B Upper Paleocene to Lower Eocene 164 A northern Sierra Nevada auriferous gravels Colton Formation 115 94 (backarc strata, Uinta Basin) (Cecil et al., 2010; Cassel et al., 2012) N=5 n=221 N=11 n=829 65 [Eocene] 96 75 95 100 C Maastrichtian to Paleocene forearc strata 115 155 N=18 n=224 B central Sierra Nevada auriferous gravels 89 (Cecil et al., 2010) 101 148 167 N=2 n=184 [Eocene] D 106 74 Campanian to Maastrichtian 96 C southern Sierra Nevada 116 166 forearc strata 148 Goler (Paleocene) and N = 14 n = 330 [Paleocene 191 Tejon (Eocene) Formations and (Lechler and Niemi, 2011) 152 Eocene] N=3 n=276 168 E Campanian backarc strata 76 (Farrer-Tuscher) 98 D Colton Formation - 173 of Gray Canyon, Uinta basin 94 [Paleocene Uinta basin (this paper) N=4 n=144 and N=5 n=221 Eocene] 95– 106 160 98 84 F Cenomanian-Turonian forearc strata E Salinian - Mojave forearc strata N=3 n=54 (Jacobson et al., 2011) N=10 n=113 97 [Eocene] 76 G Cenomanian-Turonian backarc strata 71 167 (Dakota-Ferron) of Utah foreland F 159 169 western Transverse Ranges N=4 n=70 forearc strata 245 (Jacobson et al., 2011) [Eocene] N=6 n=83

0 50 100 150 200 250 300 50 100 150 200 250 300 Age (Ma) Age (Ma)

Figure 20. Comparative age distribution curves of arc-derived Figure 21. Comparative age distribution curves of compound arc- (younger than 285 Ma) detrital zircon (DZ) subpopulations in derived (younger than 285 Ma) detrital zircon (DZ) subpopulations Upper Cretaceous and Paleogene strata of the southern California in Paleogene strata derived from the Mojave region (blue) and the forearc (blue) and the Utah foreland backarc (red). Forearc data are Sierra Nevada (red). N is the number of samples composited for from Jacobson et al. (2011); backarc data from this paper. N is the each age spectrum, and n is the number of dated DZ grains in each number of samples composited for each age spectrum, and n is composite DZ population. the number of dated grains in each composite DZ population.

assemblages (Fig. 20A) stems from incorpora- by a hypothetical California paleoriver, which of the Cordilleran magmatic arc (Fig. 1) occu- tion of samples from farther north on the Salin- may actually have been an array of related sub- pied a paleotopographic syntaxis separating the ian block, near the Sierra Nevada, than for the parallel paleodrainage courses with nearby ter- elevated Nevadaplano (DeCelles, 2004) between forearc successions (Figs. 21E, 21F) deposited mini along the southern fl ank of the Uinta Basin. the Sierra Nevada segment of the Cordilleran arc adjacent to conjoined Mojave-Salinia blocks It is not known where the southern edge of the and the backarc Sevier thrust belt on the north (19 samples with 298 grain ages composited for Uinta Basin was located because the erosional from comparable uplands to the south, here des- Fig. 20A but only 16 samples with 196 grain limit of the basin fi ll as now preserved along the ignated the Mexicoplano (Fig. 1), between the ages for Figs. 21E, 21F). line of the Book Cliffs and Roan Cliffs was not Peninsular Ranges segment of the Cordilleran its original depositional limit. No basin-margin arc and the backarc thrust belt of the Sierra Madre SUMMARY PERSPECTIVES facies are exposed along the cliff line. Oriental (Lawton et al., 2009). The Nevadaplano The course of the California paleoriver from and Mexicoplano were comparable in width and U-Pb ages of detrital zircons indicate that south to north through the Utah foreland was a geotectonic setting to the modern Altiplano of arkosic sand of the fl uviodeltaic Colton For- continuation of longitudinal sediment delivery the Andes (Dickinson, 2011). mation in the Laramide Uinta Basin was not that had persisted in varying form and volume The California paleoriver that carried Mojave derived, as once thought, from Laramide uplifts, since initiation of the Cordilleran foreland basin detritus north to the Uinta Basin south of the but from the Mojave segment of the Cordilleran in mid-Cretaceous time. Laramide deformation, Uinta uplift and the sister Idaho paleoriver that magmatic arc and associated Yavapai-Mazatzal however, infl uenced the confi guration of fore- later carried Challis detritus south to the Green basement of the Mogollon highlands in south- land paleodrainages by erecting uplifts as local River Basin north of the Uinta uplift (Chetel west Laurentia. Delivery of Mojave-derived barriers to fl uvial paleofl ow. The headwaters of et al., 2011) were complementary longitudinal detritus to the Utah foreland was accomplished the California paleoriver in the Mojave segment drainages that dominated Laramide sediment

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dispersal west of the major Laramide uplifts Andersen, D.W., and Picard, M.D., 1974, Evolution of syn- Dickinson, W.R., 2011, The place of the Great Basin in the forming the heart of the Laramide province in orogenic clastic deposits in the intermontane Uinta Cordilleran orogen, in Steininger, R., and Pennell, B., basin of Utah, in Dickinson, W.R., ed., Tectonics and eds., Great Basin evolution and metallogeny: Reno, Colorado and Wyoming. The recognition of sedimentation: Society of Economic Paleontologists Geological Society of Nevada 2010 Symposium, these complementary drainages is a major step and Mineralogists Special Publication 22, p. 1167–189. p. 419–436. Andersen, T., 2005, Detrital zircons as tracers of sedimen- Dickinson, W.R., and Gehrels, G.E., 2008, Sediment deliv- forward in understanding the paleogeomorphol- tary provenance: Limiting conditions from statistics ery to the Cordilleran foreland basin: Insights from ogy of the western United States. and numerical simulation: Chemical Geology, v. 216, U-Pb ages of detrital zircons in Upper Jurassic and p. 249–270, doi:10.1016/j.chemgeo.2004.11.013. Cretaceous strata of the Colorado Plateau: American Anderson, L.P., and Dinter, D.A., 2010, Deformation and Journal of Science, v. 308, p. 1041–1082, doi:10.2475 DISCUSSION sedimentation in the southern Sevier foreland, Red /10.2008.01. Hills, southwestern Utah, in Carney, S.M., et al., eds., Dickinson, W.R., and Gehrels, G.E., 2009, U-Pb ages of Our analysis of Colton provenance and sedi- Geology of south-central Utah: Utah Geological Asso- detrital zircons in Jurassic eolian and associated sand- ciation Publication 39, p. 194–224. stones of the Colorado Plateau: Evidence for trans conti- ment dispersal illustrates how U-Pb ages for Aschoff, J., and Steel, R., 2011, Anomalous clastic wedge nental dispersal and intraregional recycling of sediment: detrital zircons can be combined with infor- development during the Sevier-Laramide transition, Geological Society of America Bulletin, v. 121, p. 408– North American Cordilleran foreland basin, USA: Geo- 433, doi:10.1130/B26406.1. mation about petrofacies and paleocurrents to logical Society of America Bulletin, v. 123, p. 1822– Dickinson, W.R., and Gehrels, G.E., 2010a, Insights into develop integrated interpretations for the origin 1835, doi:10.1130/B30248.1. North American paleogeography and paleotectonics of sedimentary assemblages. We show further Bowers, W.E., 1972, The Canaan Peak, Pine Hollow, and from U-Pb ages of detrital zircons in Mesozoic strata Wasatch Formations in the Table Cliff region, Utah: of the Colorado Plateau: International Journal of Earth that the subdivision of detrital zircon popula- U.S. Geological Survey Bulletin 1331–B, p. B1–B39. Sciences, v. 99, p. 1247–1265, doi:10.1007/s00531 tions into constituent subpopulations for statisti- Bowers, W.E., 1991, Geologic map of Bryce Canyon -009-0462-0. cal analysis can lead to improved understanding National Park and vicinity, southwestern Utah: U.S. Dickinson, W.R., and Gehrels, G.E., 2010b, Synoptic record Geological Survey Miscellaneous Investigations Series in space and time of provenance relations for Mesozoic of sediment mixing from sources contributing Map I–2108, scale 1:24,000, 15 p. strata in south-central Utah from U-Pb ages of detrital zircons of different ages. Net U-Pb age spectra Cashion, W.B., 1967, Geology and fuel resources of the zircons, in Carney, S.M., et al., eds., Geology of south- Green River Formation, southeastern Uinta basin, Utah central Utah: Utah Geological Association Publication for individual samples commonly merge age and Colorado: U.S. Geological Survey Professional 39, p. 178–193. spectra that are signals of different provenances, Paper 548, 48 p. Dickinson, W.R., Ingersoll, R.V., and Graham, S.A., 1979, which served in combination to deliver sedi- Cashion, W.B., 1973, Geologic and structure map of the Paleogene sediment dispersal and paleotectonics in Grand Junction quadrangle, Colorado and Utah: U.S. northern California: Geological Society of America ment to depositional sites by different dispersal Geological Survey Miscellaneous Investigations Series Bulletin, v. 90, p. 897–898, 1458–1528, doi:10.1130 pathways that shifted over time. Our approach Map I-736, scale 1:250,000. /GSAB-P2-90-1458. enhances the value of detrital zircons for prov- Cassel, E.J., Grove, M., and Graham, S.A., 2012, Eocene Dickinson, W.R., Lawton, T.F., and Inman, K.F., 1986, drainage evolution and erosion of the Sierra Nevada Sandstone detrital modes, central Utah foreland region: enance analysis. batholiths across northern California and Nevada: Stratigraphic record of Cretaceous–Paleogene tectonic American Journal of Science, v. 312, p. 117–144. evolution: Journal of Sedimentary Petrology, v. 56, ACKNOWLEDGMENTS Cather, S.M., Connell, S.D., Chamberlin, R.M., McIntosh, p. 276–293. W.C., Jones, G.E., Potochnik, A.R., Lucas, S.G., and Dickinson, W.R., Lawton, T.F., and Gehrels, G.E., 2009, Pat Abbott called our attention to the paleogeo- Johnson, P.S., 2008, The Chuska erg: Paleogeomor- Recycling detrital zircons: A case study from the Cre- graphic constraints imposed by Paleogene paleo- phic and paleoclimatic implications of an Oligocene taceous Bisbee Group of southern Arizona: Geology, sand sea on the Colorado Plateau: Geological Society v. 37, p. 503–506, doi:10.1130/G25646A.1. drainages delineated by him and his colleagues in of America Bulletin, v. 120, p. 13–33, doi:10.1130 Dickinson, W.R., Cather, S.M., and Gehrels, G.E., 2010, southernmost California and adjacent Baja California. /B26081.1. Detrital zircon evidence for derivation of arkosic sand We appreciate discussions of Laramide lacustrine Cecil, M.R., Ducea, M.N., Reiners, P., Gehrels, G., Mulch, in the eolian Narbona Pass Member of the Eocene- sedimentation with Alan Carroll. We thank Elizabeth A., Allen, C., and Campbell, I., 2010, Provenance of Oligo cene Chuska Sandstone from Precambrian base- Cassel, Robinson Cecil, Gary Hunt, Carl Jacobson, Eocene river sediments from the central northern Sierra ment rocks in central Arizona, in Fasssett, J.E., et al., and Paul Link for providing Excel fi les of U-Pb ages Nevada and implications for paleotopography: Tec- eds., Geology of the Four Corners country: New Mex- for their detrital zircon samples, and Tom Fouch for tonics, v. 29, TC6010, doi:10.029/2010TC002717, 13 p. ico Geological Society 61st Annual Field Conference a detailed measured section of the strata exposed in Chan, M.A., and Pfaff, B.J., 1991, Fluvial sedimentology Guidebook, p. 125–134. of the Upper Cretaceous Castlegate Sandstone, Book Dickinson, W.R., Lawton, T.F., Pecha, M., and Gehrels, Gate Canyon. Paul Heller provided data from the Cliffs, Utah, in Chidsey, T.C., Jr., ed., Geology of east- G.E., 2011, Provenance of Paleogene Colton Formation unpublished thesis of Genevive Mathers. Brady Fore- central Utah: Utah Geological Association Publication (Uinta basin, NE Utah): Constraints from combined man discussed with us his unpublished paleocurrent 19, p. 95–109. petrofacies and DZ analysis: Geological Society of data for the DeBeque Formation. Carol A. Hill pro- Chapin, C.E., 2012, Origin of the Colorado Mineral Belt: America Abstracts with Programs, v. 43, no. 4, p. 1–2. vided guidance to Duff Brown Tank (sample COL13). Geosphere, v. 8, p. 28–43, doi:10.1130/GES00694.1. Doelling, H.H., and Davis, F.D., 1989, The geology of Kane Comments by Jon Spencer improved the text. 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