Sediment
Sediment comprises solid particles, ranging from clay-sized grains to boulders, derived primarily from the mechanical and chemical weathering of preexisting rocks, as well as from biological sources such as shell fragments or plant debris, that are transported by natural agents including water, wind, and ice before settling in depositional environments.[1][2][3] These particles undergo classification into clastic sediments, formed from eroded rock fragments; chemical sediments, precipitated directly from aqueous solutions; and biochemical or organic sediments, resulting from the accumulation of biological remains like calcium carbonate shells or silica tests.[4][3] The formation process begins with weathering, which breaks down source materials, followed by erosion and transport that sort particles by size and density, culminating in deposition where reduced energy allows settling, and eventual lithification through compaction and cementation to form sedimentary rocks.[3][5] Sediments play a critical role in Earth science by preserving stratigraphic records of ancient environments, climates, and life forms through fossils and geochemical signatures, while also influencing modern landscapes, nutrient cycling, and coastal resilience via processes like delta formation and barrier island maintenance.[6][7][8] In aquatic systems, excessive sediment loads can degrade habitats by reducing light penetration and smothering benthic organisms, underscoring the balance between natural deposition and anthropogenic impacts.[9][10]Definition and Origins
Definition and Fundamental Characteristics
![Sediment plume in sea showing suspended particles][float-right]Sediment consists of solid particulate matter derived from the weathering and erosion of pre-existing rocks, minerals, or organic materials, which is transported by agents such as water, wind, ice, or gravity and subsequently deposited in a new location.[2] This material ranges in origin from inorganic fragments like quartz grains and clay minerals to biogenic remains such as shells and plant debris, accumulating as loose, unconsolidated layers in sedimentary basins, riverbeds, ocean floors, or deserts.[1] Unlike sedimentary rocks, which form through lithification processes involving compaction and cementation, sediment exists in a granular, non-cohesive state prior to diagenesis.[11] Key fundamental characteristics of sediment include variability in particle size, spanning from submicroscopic clay (<0.002 mm) to large boulders (>256 mm), which influences its transportability and depositional behavior.[2] Particles are typically classified by diameter into categories such as gravel, sand, silt, and clay, with finer grains settling more slowly from suspension and coarser ones requiring higher energy for entrainment.[12] Compositionally, sediments are predominantly clastic (detrital), formed from mechanical breakdown of source rocks, but may also include chemical precipitates like evaporites or biochemical accumulations such as carbonates from marine organisms.[13] The unconsolidated texture allows for resuspension under sufficient shear stress, reflecting a dynamic equilibrium between erosion, transport, and deposition governed by fluid dynamics and gravitational forces.[9] Sediments exhibit primary depositional fabrics, such as bedding or lamination, resulting from sequential settling under varying flow conditions, which preserve records of paleoenvironments and transport history.[3] Organic content, often comprising less than 5% by volume in clastic sediments but higher in biogenic types, contributes to geochemical properties like porosity and permeability, affecting water flow and nutrient cycling in aquatic systems.[9] These characteristics underscore sediment's role as a transient phase in the rock cycle, linking weathering at source areas to eventual rock formation through burial and transformation.[14]
Sources and Generation Processes
Sediment originates mainly from the mechanical and chemical breakdown of pre-existing rocks, which can be igneous, metamorphic, or sedimentary in nature, through weathering processes that produce loose particles ranging from clay to boulders.[5] Additional sources include biological activity, such as the accumulation of skeletal remains from marine organisms like foraminifera and mollusks, and chemical precipitation from supersaturated solutions in aquatic environments.[15] Volcanic eruptions contribute pyroclastic materials, including ash and lapilli, while extraterrestrial inputs like cosmic dust represent minor sources, typically less than 1% of global sediment flux.[16] Physical weathering generates clastic sediments by disintegrating bedrock without changing its mineral composition, driven by mechanisms such as frost wedging—where water expands upon freezing in cracks, exerting pressures up to 30 MPa—and thermal expansion from diurnal temperature fluctuations in arid regions, which can fracture rocks along planes of weakness.[17] Exfoliation, observed in granite domes like those in Yosemite National Park, occurs due to unloading of overlying material, reducing confining pressure and causing outer layers to peel away in sheets up to several meters thick.[18] These processes dominate in cold or temperate climates and on steep slopes, producing angular fragments that retain the source rock's grain size distribution. Chemical weathering produces finer sediments by altering primary minerals into secondary ones through reactions with atmospheric gases, water, and ions, with hydrolysis of feldspars to clays being a primary pathway that accounts for much of the global clay sediment budget.[3] Oxidation of iron-bearing minerals forms rust-colored residues, as seen in the weathering of basalts to laterites, while dissolution of carbonates by carbonic acid—generated from CO2 dissolution in rainwater, reaching pH levels around 5.6—yields calcium bicarbonate solutions that can precipitate as calcite elsewhere.[16] Rates vary; for instance, in tropical environments with high rainfall (over 2000 mm annually), chemical weathering can remove up to 0.1 mm of rock per year, compared to negligible rates in dry deserts.[17] Biogenic processes contribute organic and bioclastic sediments, where shells, tests, and plant debris accumulate; for example, coral reefs generate aragonite sands from fragmented skeletons, comprising up to 90% of nearshore sediments in tropical settings.[15] In continental settings, peat from undecayed vegetation in wetlands can compact into lignite, with global organic carbon burial rates estimated at 0.1-0.2 GtC per year.[3] Chemical sediments form via inorganic precipitation, such as evaporites like halite in restricted basins where evaporation exceeds inflow, or iron-manganese nodules on ocean floors via oxidation of dissolved metals at redox boundaries.[18] Erosion then mobilizes these generated particles, with detachment thresholds influenced by shear stress from overland flow or wind, typically requiring velocities of 0.5-1 m/s for fine sands.[17]Physical Properties and Classification
Grain Size and Distribution
Grain size in sediments refers to the diameter of individual particles, typically measured in millimeters or using the logarithmic phi (φ) scale, where φ = -log₂(d) and d is the particle diameter in millimeters. This classification enables standardized description of clastic sediments, with the Udden-Wentworth scale serving as the primary framework, defining categories from clay (<0.0039 mm) to boulders (>256 mm).[19][20] The Wentworth scale, formalized in 1922, uses geometric progression for boundaries, with each coarser class approximately double the size of the finer one, facilitating analysis of transport and depositional processes.[21]| Category | Subcategory | Diameter (mm) |
|---|---|---|
| Gravel | Boulder | >256 |
| Cobble | 64–256 | |
| Pebble | 4–64 | |
| Granule | 2–4 | |
| Sand | Coarse | 0.5–1 |
| Medium | 0.25–0.5 | |
| Fine | 0.125–0.25 | |
| Very fine | 0.0625–0.125 | |
| Silt | - | 0.0039–0.0625 |
| Clay | - | <0.0039 |
Particle Shape and Texture
Sediment particle shape encompasses sphericity, which measures the degree to which a grain approximates a sphere, and roundness, which quantifies the sharpness of its edges and corners. Sphericity is defined as the ratio of the surface area of a sphere with the same volume as the particle to the actual surface area of the particle, yielding values from 0 to 1, where 1 indicates a perfect sphere.[28] Roundness, independent of sphericity, is often assessed via the ratio of the average radius of curvature of grain corners to the radius of the maximum inscribed circle, with angular particles having low values and well-rounded ones approaching 1.[29] These properties are influenced by the grain's source material and subsequent transport, where initial shapes derive from bedrock fracturing, often producing elongated or platy forms in crystalline rocks, while metamorphic sources yield more equant grains.[30] Quantitative analysis of shape employs methods like the Zingg classification, which categorizes particles based on ratios of their three principal axes (longest, intermediate, shortest), distinguishing spheres, rods, discs, and blades.[31] Modern techniques, including dynamic image analysis and Fourier descriptors, enable precise computation of sphericity and roundness from digital images, revealing that fluvial transport progressively increases both due to abrasion, with aeolian processes enhancing roundness through impacts but preserving some angularity in coarse fractions.[28] Glacial transport, by contrast, minimally alters shape, resulting in angular, low-sphericity grains from mechanical plucking and limited sorting. Particle texture refers to microscale surface features, discernible via scanning electron microscopy (SEM), which record the history of mechanical, chemical, and biological interactions. Common textures include conchoidal fractures and percussion marks from high-energy collisions in fluvial or beach environments, upturned plates from glacial crushing, and chemical dissolution pits in low-energy or acidic settings.[32] SEM studies classify quartz grain surfaces into categories such as abraded (smoothed by collision), corroded (etched by solution), and frosted (dull from wind abrasion), with beach sands showing high frequencies of abrasion features and eolian sands exhibiting delicate frosting absent in glacial deposits.[33] These textures provide diagnostic evidence of depositional environments; for instance, blocky, faceted grains indicate subglacial comminution, while smooth, featureless surfaces suggest prolonged chemical weathering.[34] Surface roughness, quantified by fractal dimensions or Fourier coefficients, correlates with transport duration, decreasing as grains polish through repeated impacts.[35]Composition and Mineralogical Content
Sediments consist of detrital grains derived from the physical and chemical breakdown of bedrock, supplemented by biogenic remains and chemical precipitates, with mineralogical composition reflecting the parent rock lithology, weathering intensity, and diagenetic alterations. Siliciclastic sediments, the most widespread type, are dominated by silicate minerals such as quartz (SiO₂), which constitutes the primary framework component due to its high resistance to chemical weathering and mechanical abrasion during transport.[36] Accessory framework minerals typically include feldspars (e.g., orthoclase, plagioclase, microcline), micas (e.g., biotite, muscovite), and lithic fragments, alongside trace heavy minerals like zircon, rutile, and tourmaline, which persist as stable remnants in mature sediments.[37] [38] In finer-grained siliciclastic fractions, particularly muds and clays, phyllosilicate minerals predominate, including illite, smectite, chlorite, and kaolinite, formed through hydrolysis and neoformation during weathering of feldspars and mafic silicates. These clay minerals exhibit variable abundances influenced by climate and provenance; for instance, smectite is enriched in sediments from volcanic or arid source regions due to incomplete weathering, while kaolinite prevails in humid, intensely leached environments.[39] Heavy minerals such as magnetite, ilmenite, and garnet may also occur, providing provenance indicators, though their concentrations rarely exceed a few percent. Cementing phases in partially consolidated sediments often comprise authigenic quartz, calcite (CaCO₃), or iron oxides like hematite (Fe₂O₃), binding grains post-deposition.[40] Carbonate sediments, by contrast, are chiefly composed of biogenic calcite and aragonite (CaCO₃ polymorphs) from skeletal debris of marine organisms such as foraminifera, mollusks, and corals, with lesser siliciclastic admixtures in mixed systems. Dolomite (CaMg(CO₃)₂) can form diagenetically via magnesium enrichment in interstitial waters, altering primary aragonite. Biogenic siliceous sediments, including diatomaceous oozes and radiolarian cherts, feature opal (hydrated SiO₂) as the key mineral, while evaporitic sediments incorporate halides like halite (NaCl) and sulfates such as gypsum (CaSO₄·2H₂O). Organic-rich sediments may include kerogen precursors or phosphates from biogenic sources, but mineralogically, they align with associated clastic or carbonate matrices.[3] Overall, mineral assemblages in sediments serve as proxies for tectonic setting and paleoclimate, with quartz-feldspar ratios indicating immature, proximal sources versus quartz-rich, recycled distal ones.[38]Transport Mechanisms
Initiating Forces and Thresholds
The initiation of sediment transport requires the application of forces that overcome the resistive forces acting on particles, such as submerged weight, interparticle friction, and cohesion. In fluid-dominated environments, the primary initiating forces are hydrodynamic: drag, which acts parallel to the flow and results from pressure differences around the particle, and lift, a perpendicular force arising from turbulent eddies and boundary layer effects that reduce effective particle weight. These forces are generated by shear stress at the bed interface, proportional to fluid density, velocity squared, and particle projected area.[41][42] In slope-dominated settings, the tangential component of gravity provides an additional force, enhanced by oversteepening or seismic triggers, leading to detachment via sliding or rolling. For non-cohesive particles, resistance is primarily geometric and density-dependent, while cohesive fines (silt and clay) exhibit higher thresholds due to electrostatic and van der Waals bonding, often requiring 10-100 times greater shear stress than predicted for non-cohesive equivalents.[43][44] The threshold for particle entrainment, or incipient motion, is defined as the minimum condition under which dislodgement occurs more frequently than re-embedding, typically expressed as critical bed shear stress (τ_c). This threshold varies with particle size (d), density contrast, fluid viscosity, and bed roughness; for gravel and sand in water, τ_c increases with d up to ~1-2 mm before stabilizing due to inertial dominance. The dimensionless Shields parameter encapsulates this: θ_c = τ_c / [(ρ_s - ρ_f) g d], where ρ_s is sediment density (~2650 kg/m³ for quartz), ρ_f is fluid density, and g is gravitational acceleration (9.81 m/s²). Experimental data from flume studies indicate θ_c ≈ 0.045-0.06 for turbulent flows over flat, non-cohesive beds with Reynolds numbers >500, though values drop to ~0.03 on rippled beds or rise above 0.1 for armored gravel surfaces.[41][42][43] Empirical curves like the Hjulström diagram relate critical flow velocity (u_c) to grain size, showing a U-shaped entrainment curve: u_c decreases from ~100 cm/s for coarse gravel (d > 10 mm) to a minimum of ~20-30 cm/s for medium sand (0.2-0.5 mm), then rises sharply for silt (<0.06 mm) due to cohesion, exceeding 50 cm/s for particles <0.01 mm. In aeolian settings, thresholds are lower owing to the fluid's lower density; critical friction velocity (u_*c) for sand is ~0.2-0.3 m/s, scaling with sqrt(θ_c (ρ_s - ρ_air)/ρ_air). Field measurements from rivers confirm these thresholds are rarely exceeded; for instance, in the Yellowstone River (2005 study), bedload initiation for 2-8 mm gravel required velocities >1 m/s, occurring <5% of the time under median flows. Biological and packing effects can elevate thresholds by 20-50%, as seen in vegetated or sorted beds.[41][45][46]| Grain Size Class | Typical Critical Velocity (cm/s, water) | Shields Parameter Range | Notes |
|---|---|---|---|
| Clay (<0.002 mm) | >100 (cohesive) | >0.1 | Cohesion dominates; lab flume data.[41] |
| Silt (0.002-0.06 mm) | 50-70 | 0.05-0.1 | Higher due to partial cohesion.[46] |
| Sand (0.06-2 mm) | 20-40 | 0.03-0.06 | Minimum entrainment; turbulent flow.[42] |
| Gravel (2-64 mm) | 30-100+ | 0.04-0.05 | Increases with size; field/river data.[45] |
Fluvial and Riverine Transport
Fluvial sediment transport encompasses the movement of particulate material within rivers and streams, primarily driven by water flow exceeding the critical shear stress for particle entrainment.[47] This process is divided into bedload and suspended load modes, with bedload involving coarser particles (typically sand and gravel) that roll, slide, or saltate along the channel bed, maintaining frequent contact with the substrate.[48] Saltation occurs as particles are lifted briefly by turbulent bursts before settling downstream, contributing to bedload flux that can reshape channel morphology over time.[49] Suspended load, comprising finer silt and clay particles, is upheld in the water column by upward turbulent velocities that counteract settling.[10] These particles follow flow streamlines with minimal bed interaction, often dominating total sediment discharge in rivers with high turbidity; for example, one analysis found suspended load accounting for 93.5% of transport in a studied watershed.[50] Wash load, a subset of suspended material too fine to deposit under typical flows, originates from distant upstream sources and remains perpetually aloft. Quantitative prediction of bedload relies on empirical formulas like the Meyer-Peter and Müller equation, formulated in 1948 from gravel-bed flume data, which expresses transport rate as proportional to the cube root of excess boundary shear stress beyond the critical threshold.[51] The formula, \phi = 8(\theta - \theta_c)^{3/2} where \phi is the dimensionless transport rate and \theta the Shields parameter, assumes uniform non-cohesive sediment and steady uniform flow, though extensions address variability.[52] Suspended load estimation often couples advection-diffusion models with settling velocities, calibrated against discharge-sediment rating curves derived from gauging station data.[53] Transport efficiency varies with hydraulic parameters: higher velocities and discharges during floods mobilize larger grains and increase suspension heights, while channel slope and sediment supply dictate equilibrium profiles.[54] In gravel-bed rivers, bedload forms riffles and pools through differential transport, whereas sand-bed systems favor plane-bed configurations with prevalent suspension.[55] Upstream supply limitations, such as from dams, can reduce downstream flux by over 90% in regulated basins, altering habitats and delta progradation.[47] Deposition predominates in velocity-reduced zones like meander bends or overbank floods, where reduced shear promotes settling and bar formation.[56]Aeolian and Wind-Driven Transport
Aeolian transport refers to the movement of sediment particles by wind, predominantly in environments with low vegetation cover such as deserts, beaches, and periglacial regions. This process encompasses entrainment, where particles are lifted from the surface; transport via saltation, suspension, or creep; and eventual deposition when wind energy diminishes. Entrainment begins when wind-generated shear stress surpasses the threshold friction velocity, typically around 0.2-0.3 m/s for dry, loose quartz sand grains of 0.1-0.5 mm diameter, equivalent to a near-surface wind speed of approximately 4-6 m/s.[57][58] The primary modes of transport are saltation, suspension, and surface creep. Saltation dominates for medium to coarse sand (0.06-2 mm), involving grains ejected to heights of up to 2 m and horizontal distances of centimeters to meters via ballistic trajectories influenced by drag and gravity; upon impact, these grains may dislodge others, sustaining transport. Suspension applies to finer silt and clay particles (<0.06 mm), which remain airborne for extended periods and can travel hundreds of kilometers, contributing to dust storms and loess deposits. Surface creep, or reptation, affects larger grains (>0.5 mm) or smaller particles propelled by saltating impacts, resulting in rolling or sliding along the bed at rates comprising 5-25% of total sand flux.[58][59][60] Transport rates are quantified by empirical relations such as Bagnold's formula, which scales flux proportional to the cube of excess shear velocity and inversely with grain density and size, emphasizing wind speed as the key driver. Factors modulating aeolian transport include moisture content, which elevates thresholds by 2-10 times via surface tension; vegetation and roughness elements that reduce effective wind shear; grain size distribution, with optimal transport for unimodal fine sands; and fetch distance, where longer upwind expanses allow flux equilibration. In coastal settings, wave-driven wetting and drying cycles further influence availability, while atmospheric turbulence can lower entrainment thresholds by enhancing lift forces.[57][61]Glacial, Gravity, and Mass Wasting Transport
Glaciers entrain and transport sediment primarily through subglacial, englacial, and supraglacial pathways, where ice deformation and meltwater play key roles in movement. Subglacial transport occurs via basal plucking, where ice freezes to bedrock irregularities and uplifts fragments, or through abrasion that grinds rock into finer particles; these sediments are then carried forward by glacier sliding or bed deformation, often forming thick debris layers in marginal zones. Englacial transport involves debris embedded within the ice mass during flow, while supraglacial transport handles surface debris from rockfalls or supraglacial streams, which can be redistributed via melting. Active subglacial pathways dominate in temperate glaciers, moving larger volumes compared to passive supraglacial routes, as documented in models distinguishing high-efficiency basal entrainment from slower surface accumulation.[62][63] Meltwater streams emerging from glaciers further enhance transport by suspending fine lithogenic particles eroded from bedrock, generating high sediment loads in proglacial environments; for instance, turbidity currents driven by meltwater have been observed to redistribute glacially derived material across fjords and plains. In deforming-bed glaciers, shear within the sediment layer itself facilitates bulk transport, with rates potentially exceeding those of rigid-bed systems by factors of 10 or more in soft substrates.[64] Gravity-driven transport, distinct from fluid-mediated processes, occurs on slopes where sediment moves downslope primarily under gravitational pull without dominant water or wind influence, often manifesting as slow creep or colluvial accumulation. Colluvium forms as heterogeneous, poorly sorted deposits—containing less than 50% material larger than 60 mm—transported by sheetwash, rainwash, or dry mass gravity flows on gentle to moderate slopes, resulting in unstratified aprons at slope bases. These processes dominate in arid or periglacial settings, where freeze-thaw cycles or solifluction mobilize regolith, contributing to hillslope sediment budgets before fluvial interception.[65] Mass wasting encompasses rapid gravity-dominated downslope movements of unconsolidated sediment and rock, triggered when slope stability thresholds are exceeded by factors like oversteepening, water saturation, or seismic activity. Key types include rockfalls (free-falling blocks), rotational slides (coherent slump blocks on curved failure planes), translational slides (planar gliding), and flows such as debris flows, where saturated sediment mixtures behave as viscous slurries with 40-70% solids by volume, capable of entraining boulders up to several meters in diameter. Earthflows and mudflows represent finer-grained variants, with velocities ranging from centimeters per day in creeps to over 10 m/s in catastrophic debris avalanches, eroding and redepositing sediment in talus cones or aprons. These events supply significant coarse fractions to drainage systems, with historical examples like the 1980 Mount St. Helens eruption mobilizing over 2 billion cubic meters of sediment via mass flows. Fluid presence facilitates but does not dominate, as gravity provides the primary driving force on slopes exceeding 5-10 degrees.[66][67][68]Depositional Environments
Terrestrial and Continental Settings
Terrestrial and continental settings encompass non-marine depositional environments where sediments accumulate through fluvial, aeolian, lacustrine, and glacial processes, primarily in river valleys, deserts, lakes, and glaciated regions. These environments produce distinct sediment characteristics driven by transport mechanisms and local topography, such as fining-upward sequences in fluvial systems and unsorted mixtures in glacial tills. Sedimentation rates vary widely, from millimeters per year in stable floodplains to meters per year in proglacial zones, influenced by climate, relief, and vegetation cover.[69] In fluvial and alluvial settings, rivers deposit coarse gravels and sands in channel beds and point bars during high-flow events, transitioning to finer silts and clays on floodplains as velocity decreases. These deposits exhibit cross-bedding, scours, and fining-upward cycles, reflecting episodic flooding and waning flows; for instance, braided rivers form sheet-like gravel sheets, while meandering rivers build levees and overbank fines. Alluvial fans at mountain fronts spread coarse debris radially, with grain size decreasing downslope due to reduced competence. Such sediments cover approximately 23% of ice-free continental areas globally.[70][71][72] Lacustrine environments yield fine-grained, laminated muds and clays from suspended load settling in standing water, often with varves—annual layers of silt and clay—recording seasonal variations in inflow. Organic-rich deposits dominate in productive lakes, while coarser deltas form at inflows; these sediments are typically well-sorted and low-energy indicators, with thicknesses reaching hundreds of meters in tectonic basins like ancient Lake Bonneville. Deposition occurs at rates of 0.1–10 mm/year, preserving delicate structures due to minimal post-depositional disturbance.[73][74] Aeolian deposits in arid continental interiors consist of well-sorted, rounded quartz sands forming dunes and sheets, transported by saltation and suspension. Loess, fine silt from glacial outwash, blankets vast areas like the Chinese Loess Plateau, with thicknesses exceeding 300 meters, deposited by prevailing winds at rates up to 0.5 mm/year. These sediments show frosted grains, high sphericity, and deflation hollows, covering about 21% of continental surfaces.[75][72][76] Glacial continental deposits include lodgement till—compacted, unsorted mixtures of clay to boulders—directly emplaced subglacially, and meltout till from supraglacial debris, forming ground moraines and drumlins. These cover roughly 20% of ice-free land, with particle sizes spanning five orders of magnitude and fabrics aligned by ice flow; eskers and kames represent sorted glaciofluvial infills in subglacial channels. Till sheets can exceed 100 meters thick in continental ice sheets like Laurentide.[77][78][72]Marine and Oceanic Basins
![Sediment distribution in the Gulf of Mexico][float-right]Marine and oceanic basins encompass a range of depositional environments, from continental shelves and slopes to abyssal plains, where sediments accumulate through settling of fine particles, biogenic remains, and episodic gravity flows. On continental shelves, neritic sediments predominate, consisting primarily of terrigenous sands and silts derived from coastal erosion and river inputs, with deposition influenced by waves and currents that sort grains by size. These areas cover approximately 25% of the seafloor and feature higher accumulation rates compared to deeper settings, often exceeding 1 cm per 1,000 years in proximal zones. [79][80] In deeper oceanic basins, pelagic sedimentation dominates, characterized by slow settling of fine-grained terrigenous clays, biogenic oozes, and minor hydrogenous components through the water column. Pelagic sediments, which blanket about 75% of the ocean floor, accumulate at rates typically ranging from 0.1 to 10 mm per 1,000 years, reflecting the vast distances from land sources and limited supply of coarse material. Calcareous oozes, formed from the tests of planktonic foraminifera and coccolithophores, prevail in shallower basins above the carbonate compensation depth (around 4,000-5,000 meters), comprising over 30% biogenic calcium carbonate, while siliceous oozes from diatoms and radiolarians occur in nutrient-rich upwelling zones or below the CCD. [79][80][81] Abyssal plains in oceanic basins receive hemipelagic and pelagic fines, including red clays in areas starved of biogenic input, where aluminum-rich clays from atmospheric dust and volcanic ash settle uniformly. Episodic turbidity currents transport coarser terrigenous sands and silts via submarine canyons to basin floors, forming submarine fans and turbidite sequences characterized by graded bedding in Bouma cycles, with individual event beds up to several meters thick. These fans, such as those in the Gulf of Mexico or Bengal Fan, can accumulate thicknesses exceeding 10 km over geological time, driven by density flows that bypass shelves during sea-level lowstands or storms. [82][81][83] Sediment distribution in basins reflects bathymetric controls and ocean circulation, with trenches and fracture zones trapping additional material from subducting plates, leading to localized thickening. Biogenic productivity gradients dictate ooze types, with calcareous forms covering roughly 48% of the global seafloor and siliceous about 4%, while clays fill the remainder in low-productivity regions. [84]