Abstract: A well-preserved record of phosphatic sedimentation across an ancient continental ramp is contained in Upper Leonardian–Guadalupian (Upper Permian) strata of the Phosphoria and Park City Formations exposed along the flanks of the Uinta Mountains, northeastern Utah. The Uinta Mountain exposures trend approximately down depositional dip and, unlike exposures of the type Phosphoria in the Idaho–Wyoming overthrust belt, are not structurally telescoped. This study is based on 12 detailed measured sections and hundreds of samples collected and studied macroscopically and microscopically from the uppermost Grandeur Member (Park City Formation), Meade Peak Member (Phosphoria Formation), and Franson Member (Park City Formation). The uppermost Grandeur Member is characterized by local carbonate hardgrounds, which are partially reworked in a transgressive lag at the base of the Meade Peak. Strata from the Meade Peak and Franson are described in terms of five lithofacies: 1) organic-rich dolomite, which is fine-grained, laminated, and present only in distal sections of the Meade Peak; 2) granular phosphorite packstone and grainstone, which is mainly associated with the Meade Peak and consists of peloids, phosphatic shelly debris, intraclasts, molluscan steinkerns, and local phosphatic coated grains; 3) gray shale; and 4) chert, which are most common at the transition zone between the Meade Peak and Franson Members; and 5) phosphatic dolowackestone, which comprises most of the Franson and which contains calcite-filled molds after anhydrite nodules and supratidal fabrics in more shoreward sections. Stratal architecture and stacking patterns, as well as the sedimentologic character and degree of bioturbation of individual facies, suggest that the Meade Peak and Franson members together are unconformity bounded and record a single major transgressive-regressive cycle. As such, they are interpreted as a stratigraphic sequence. In addition, detailed sedimentologic relations suggest that repetitively interbedded phosphorite, gray shale, and phosphatic dolowackestone of the transition zone between the Meade Peak and Franson were deposited in response to small-scale fluctuations in sea level, likely with associated changes in climate that affected the ramp. The nature and style of phosphate accumulation were not uniform across the ramp, but varied down depositional dip. Poorly oxygenated basin–ward positions were marked by deposition of organic-rich carbonate in which phosphate is interpreted to have grown authigenically in situ . In contrast, the inner ramp was current-winnowed, well oxygenated, and was the site of phosphate condensation, as suggested by the presence of only minor amounts of interstitial mud, the ubiquity of bioturbation, and the abundance of granular phosphorite, respectively.

Pore systems in the Middle Permian Phosphoria Rock Complex (PRC), Rocky Mountain Region, USA, evolved with biotic and chemical dynamics in a shallow epicontinental seaway undergoing extreme environmental shifts. Biochemical responses to environmental changes directly affected pore systems and controlled diagenetic pathways through burial. Petrographic methods and spatially resolved measurements of δ 18 O in sequence stratigraphic context allow characterization of pore systems and their evolution in heterogenous biochemical sediments. Pore systems vary regionally and across systems tracts on second-order (9–10 million years [MY]) and third-order (2–5 MY) timescales. Minimal porosity occurs in transgressive mudrocks rich in organic matter (OM), phosphorites, and carbonates. Cool, acidic, low-oxygen, nutrient-rich basinal waters interacted with warm open to restricted shelfal waters in transgressions. This resulted in accumulation and microbial decay of S-rich OM, phosphatization, carbonate precipitation, silicification, as well as deposition of calcitic-biotic debris (bryozoans, brachiopods, and crinoids) and micrite. Relative to landward and highstand marine components, transgressive basinal marine carbonates and silica are δ 18 O depleted due to microbial decay of OM. Extensive cementation coupled with near-surface compaction and recrystallization of micrite occluded large portions of porosity in transgressive phosphorites and carbonates. Porosity in these rocks is dominated by interparticle and, to a lesser degree, intraparticle microporosity in microbored phosphatized and micritized grains. Phosphorites are compacted where cements are not pervasive. OM-rich sediments host minimal volumes of interparticle nanoporosity due to mechanical compaction and incursion of secondary OM (bitumen) during burial. PRC OM is S-rich, due to sulfate-reducing bacterial enrichment, and locally abundant. This drove early generation of secondary OM and inhibited OM-hosted porosity development through thermal maturation. Large volumes of porosity accumulated in highstand sediments and varied with transitions from silicisponge spicule cherts and calcitic-biota carbonates to pervasively dolomitized micritic, peloidal, aragonitic mollusk, and peritidal microbial sediments. These biochemical transitions, and ultimately pore-system evolution, were driven by interaction between oxygenated open marine waters, eolian siliciclastic debris, and increasingly restricted shelfal waters. Marine carbonate and silica δ 18 O are consistent with Middle Permian open marine waters but are enriched landward and through highstands with evaporative fractionation. This δ 18 O-enriched authigenic silica in carbonates and evaporite replacements, as well as δ 18 O enrichment through silica precipitation, suggest dolomitization and silicification were driven by evaporitic processes. In spiculitic cherts and siltstones, silicification and carbonate diagenesis resulted in small volumes of intraparticle, interparticle, and moldic porosity, as well as increased susceptibility to fracturing and associated permeability enhancement. Chalcedony in spiculites and silicified carbonates host minor volumes of porosity where moganite crystallites dissolved during hydrocarbon migration. Highstand dolomites host abundant intercrystalline, moldic, fenestral, and interparticle macroporosity and microporosity, especially in peloidal wackestones, mollusk debris, ooid grainstones, and peritidal microbialites. Dolomitization resulted in dissolution of aragonitic mollusk and ooids, cementation, and preservation of primary porosity. Porosity loss through burial in dolomites occurs through mechanical compaction, and to a lesser degree, precipitation of zoned carbonate cements that are δ 18 O depleted relative to earlier dolomite. Compaction strongly decreases intercrystalline porosity in dolomitized peloidal wackestones. Secondary OM related to hydrocarbon migration coats surfaces and fills small pore volumes, inhibiting burial cementation.

Abstract This chapter deals with the removal of minerals from sandstones and mudrocks during any stage of diagenesis by pore waters undersaturated with respect to one or more minerals. Dissolution has long been known to be important in “more soluble” rocks, especially carbonates and evaporites, but the past five decades have seen a growing understanding that it is an important process in clastic terrigenous deposits as well. It should not have been a surprise—sandstones and mudrocks can comprise many different minerals, detrital as well as authigenic, that were formed under a wide range of conditions. Subsequent eogenetic, mesogenetic and telogenetic diagenetic environments subject minerals to varied temperatures and pressures and especially to diverse water chemistries that potentially can range from nearly pure meteoric water to hypersaline brines and from acidic to alkaline. As a result, one can find instances of dissolution of almost any mineral. Widespread dissolution of common grains and cements may have a significant impact on the ultimate composition of terrigenous rocks, their porosity and their reservoir potential.

Abstract Phylum Echinodermata Subphylum Echinozoa Class Echinoidea — Late Ordovician-Recent Echinoids (sea urchins) live in normal marine environments because they with a very limited range of salinity tolerance (generally only a few ppm). They occur mainly as grazers or burrowers in sandy shelf areas or as grazers and bioeroders along rocky shorelines. They occur in deeper waters as well, extending to abyssal depths. Fossil forms are most common in normal marine, open shelf or platform deposits. Echinoids are common in both warm- and cold-water settings, although they rarely are major rock-forming organisms (i.e., they rarely exceed 10-15% of the total sediment). Modern and ancient echinoids are/were composed of moderate- to high-Mg calcite. Modern forms contain between 2 and 17 mole% Mg; the Mg content varies with generic group and increases with increasing water temperature (see Milliman, 1974 , p. 130-134, for details and citations). Echinoids, like all echinoderms at some stage in their life cycle, show pentameral (five-fold) symmetry. They have heavily calcified, globular to discoidal, hollow, endoskeletal tests (coronas) that are composed of individual sutured, interlocking or imbricated calcite plates. The calcitic coronal plates are porous and sponge-like; echinoids with rapid growth rates have spongier plates (with more holes and less calcification) than slowgrowing counterparts. Thus, slow growing, cold-water forms can be more heavily calcified than those from warmer waters ( Raup. 1958 ) In life, echinoid tests are covered with elongate, moveable spines (in some species extremely short, but in others longer than 8 cm). The spines normally detach after death and can themselves be significant sediment contributors. Generally, each plate of an echinoid behaves optically as a single, extensively perforated, calcite crystal (see comments below). Echinoid teeth, however, are polycrystalline.