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Geofluids: Sedimentary basin-scale fluid flow

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Fluid flow, like rocks and sediments, define sedimentary basins

There is a complex interweave of fluid, sediment, and rock in the life of a sedimentary basin.  Beneath Earth’s surface fluids occupy every pore, fracture, nook, and cranny (except for a skinny, unsaturated veneer at the very top). Geofluids participate in nearly all post-depositional processes: they transport dissolved mass and heat, are conscripted by virtue of their chemistry to diagenetic and metamorphic domains, and provide welcome relief for otherwise difficult tectonic processes. In this context fluid flow refers to large-scale, subsurface systems that may extend 100s of kilometres laterally and to depths of 100s or 1000s of metres.

Geofluids are rarely static. Advective fluid flow results primarily from gradients in potential energy (hydraulic gradients). Fluid flow is driven by:

  • The gravitational potential energy derived from surface topography
  • Sediment compaction
  • Tectonism
  • Convection

Molecular and mechanical diffusion, commonly derived from solute concentration gradients, also contribute to diagenetic processes such as pressure solution (stylolites), the formation of cleavage in deforming rock, and metamorphic reactions deep in the crust. In shallow groundwater systems diffusion drives the expansion of contaminant plumes beyond the limits of advective flow.

 

Permeability

Basin-scale advective flow, the kind that moves the fluid mass, requires relatively continuous permeability through large volumes of sediment and rock. This does not mean that the permeability from one rock unit to another is the same, but that there is connectivity among pore and fracture porosity throughout the basin. The actual permeability can vary by many orders of magnitude between rock units; some typical values are shown in the following table.

Fluid chemistry is coupled with fluid-rock mechanics

Subsurface fluid systems are also geochemical systems. Fluid chemistry, that governs which precipitation-dissolution reactions will take place, depends on initial composition (fresh water or sea water), the composition and temperature of the rock and sediment in which it resides, fluid flow rates, and the time spent in moving from one rock type to another (residence time). Fluid chemistry also evolves with flow through the sedimentary basin; the amount of dissolved mass generally increases with time and depth of burial.

An example of the value of knowing paleotemperature profiles in sedimentary basins. The chart shows commonly observed diagenetic reactions in relation to burial temperatures and changes in fluid composition represented here by the production of organic solvents that influence pH. The depths at which reactions begin and end will depend on the local geothermal gradient. Modified from Surdam et al. 1989.

A summary of common fluid compositions and diagenetic changes with depth and burial temperatures. The solid blue line shows evolution of organic solvents, particularly organic acids, through the oil maturation window. This window may also correspond with evolution of quartz cements, clay dehydration, and significant changes to carbonate stability. Modified from Surdam et al., 1989.

Maturation of organic matter, that is strongly temperature-dependant, can change fluid chemistry as depth and temperature increase. For example, organic acid by-products produced during maturation at temperatures of 80o to about 120oC (the oil generation window) have a profound effect on pH and carbonate stability. The changes in aqueous and hydrocarbon fluid density will, in turn, be reflected in changing buoyancy that may influence flow dynamics.

 

Fluid flow regimes

As a general rule, and one that allows for simplification of a complex problem, we divide fluid regimes into topographic, compaction, and density-temperature regimes based on the primary driving mechanisms for flow. The three regimes correlate approximately to depth, with topography-driven flow the shallowest, but there is significant overlap. For example, surface-derived meteoric fluids can extend to depths of 3-5 km, but compaction can also drive flow over the same depth range.

 

Topography driven flow

Water is recharged to meteoric systems by precipitation. Some of this water evaporates, some contributes to surface runoff, and the remainder seeps to the watertable through soil, regolith, and bedrock. Groundwater flow in aquifers and aquitards is governed by hydraulic gradients, or the difference in hydraulic head at different locations in an aquifer (relative to some datum, usually sea level). The value of hydraulic head at any point in an aquifer, expressed in units of length, is a function of the potential energy available to generate fluid flow. For most groundwater systems, this energy is generated by the gravitational potential of topography.

Meteoric flow takes place at different scales, ranging from shallow flow systems where fluid is recharged and discharged across localised ridges and valleys, to more regional systems where groundwater associated with mountainous terrain can reach depths of 5 km (McIntosh and Ferguson, 2021). Note that smaller-scale flow systems (or flow cells) tend to be nested within larger systems, a characteristic first identified by J. Toth (1963 – PDF available), an example of which is shown below. Toth’s example was developed for small drainage basins, but the principles can be applied to larger sedimentary basins.

Groundwater flow systems in a small drainage basin. The flow systems are nested according to Toth’s (1963) vision of groundwater partitioning at local, intermediate, and regional scales. Vertical scale in 100s of metres; horizontal scale in 10s to 100s of kilometres.

Groundwater flow systems in a small drainage basin. The flow systems are nested according to Toth’s (1963) vision of groundwater partitioning at local, intermediate, and regional scales. Vertical scale in 100s of metres; horizontal scale in 10s to 100s of kilometres.

The penetration of relatively fresh groundwater to these depths is identified in boreholes from samples and from electrical resistivity measurements. In Gulf Coast Basin, meteoric flushing occurs to depths of at least 2000 m. In Wanganui Basin (New Zealand) low salinity meteoric waters have been intersected in wells at 1500 m; in this basin, the topographic drive for meteoric flow was derived from terrain uplifted during the Pliocene-Pleistocene (Ricketts et al., 2004).

The residence time of groundwater in these different systems also varies by orders of magnitude: in shallow systems this is measured as days to years, whereas in regional flow systems 105 to 106 years. In general, the deeper and older the groundwater system, the more saline it becomes because the rate of diagenetic reactions generally increases with (depth-dependant) temperature.

 

Compaction-driven flow

In compacting sediment, the solid framework decreases in volume at the expense of porosity. A consequence of volume reduction is that a roughly equivalent volume of interstitial pore fluid will be driven off. How far and how fast the escaping fluid will flow depends on the permeability. If the fluid can move freely there will be little change in fluid pore pressure above the ambient hydrostatic pressure. However, if permeability is also reduced, as is commonly the case during compaction, then pore pressures will increase. If permeability is seriously reduced then pore pressures can approach lithostatic conditions. Both porosity and permeability are also affected by diagenesis that also takes place during sediment burial.

Compaction in sedimentary basins is driven by sediment load – as sedimentation proceeds, the vertical load increases. The change in porosity with depth for typical mudstone-shale and sandstone profiles is illustrated in the graphs below. Initial porosities for mudstones are as high as 60-70%, but these values decrease rapidly, almost exponentially in the first few 100 metres of burial, after which the rate of porosity loss becomes more linear.

Porosity-depth trends for data compiled from many sources. The initial stages of mudstone compaction commonly show rapid porosity loss in the upper kilometre of burial. The compaction of carbonate, and subsequent decrease in porosity is strongly dependant on early stages of cementation. Figure modified from Allen and Allen, 2005, Fig. 9.3.

Porosity-depth trends for data compiled from many sources. The initial stages of mudstone compaction commonly show rapid porosity loss in the upper kilometre of burial. The compaction of carbonate, and subsequent decrease in porosity is strongly dependent on early stages of cementation. Figure modified from Allen and Allen, 2005, Fig. 9.3.

The rate at which vertical loads increase depends on the sedimentation rate. Pore pressures in many sedimentary basins are hydrostatic to depths of 2-3 km, but commonly exceed hydrostatic conditions at greater depths. In many cases, the degree of compaction disequilibrium is caused by rapid sedimentation of low permeability mudrock or salt deposits. Gulf Coast Basin, one of the most intensely studied basins anywhere, shows these disequilibrium compaction trends over much of the basin. However, formation of diagenetic cements also takes place in conjunction with compaction, contributing to elevated pore pressures. For example, quartz cements become prolific at about 3 km burial depth in basins having average geothermal gradient (25o-30oC).

A typical pore pressure – depth curve for the Gulf Coast Basin. The transition from hydrostatic conditions to pressures approaching lithostatic values at 2.5 – 4 km, coincides approximately with quartz cementation, pH buffering by organic acids produced in the oil window, and clay dehydration, all of which have a significant effect on porosity and permeability. Modified from Bethke, 1986 – PDF available.

A typical pore pressure – depth curve for the Gulf Coast Basin. The transition from hydrostatic conditions to pressures approaching lithostatic values at 2.5 – 4 km, coincides approximately with quartz cementation, pH buffering by organic acids produced in the oil window, and clay dehydration, all of which have a significant effect on porosity and permeability. Modified from Bethke, 1986 – PDF available.

Tectonic-driven flow

The role of elevated pore pressures in promoting faulting, particularly thrust faulting, is well established (Hubbert and Rubey (1959). The elastic response to tectonically derived compressional stress is a reduction of rock volume at the expense of porosity. If escaping fluids can move freely then pore pressures remain relatively stable. However, if there are permeability barriers then pore pressures will rise above hydrostatic values. Elevated pore pressures are common in tectonically active systems such as accretionary prisms, where the primary driving forces are derived from subduction accretion of oceanic sediment and volcanic rock. Elevated pore pressures in the accretionary stack, originating from tectonic compaction, also participate in the generation of thrust faults. The faults themselves may act as conduits for fluid flow towards the sea floor where expulsion is manifested as mud diapirs and gas seeps.

Examples from wells drilled into forearc basin sediments above the Hikurangi subduction accretionary wedge (New Zealand), show pressures approaching lithostatic values at depths as shallow as 500-1000 m beneath the surface (Allis et al. 1997). Such high pore pressures at these depths cannot have been generated by sediment load alone; compression associated with thrust faulting must have played a significant role.

Pressure-depth data for wells in East Coast Basin (New Zealand) showing rapid, shallow deviations towards lithostatic conditions. The forearc basin sits atop the active Hikurangi accretionary prism. Elevated pore pressures are caused primarily by active thrusting. From Allis et al., 1997 (reference given below).

Pressure-depth data for wells in East Coast Basin (New Zealand) showing rapid, shallow deviations towards lithostatic conditions. The forearc basin sits atop the active Hikurangi accretionary prism. Elevated pore pressures are caused primarily by active thrusting. From Allis et al., 1997 (reference given below).

Convection

Convective flow is generated when heated fluids become more buoyant than their surroundings (because of changes in density).  Convection is an important process that drives mantle-derived magma plumes to shallow lithospheric levels. In sedimentary basins, large-scale convective fluid flow is probably subordinate to topography, compaction, and tectonic-driven flow because of permeability anisotropies caused by dramatic variations in sediment lithology. However, buoyancy, in addition to compaction fluid drive, does play an important role during migration of hydrocarbons. Hydrocarbon traps illustrate these effects nicely, wherein the boundaries between saline water, the overlying oil leg, and gas cap are clearly demarked on resistivity and density logs. Meteoric waters heated by magma intrusion may also be incorporated into local convection cells.

 

Reference

Allis R.G, Funnell R, Zhan X. 1997. Fluid pressure trends in basins astride New Zealand plate boundary zone, in Geofluids II Extended Abstracts, pp. 214 –217, ed., Hendry J.P. et al., Queens Univ. Belfast.

 

Other posts in this series

Sedimentary basins: Regions of prolonged subsidence

Defining the lithosphere

The rheology of the lithosphere

Isostasy: A lithospheric balancing act

The thermal structure of the lithosphere

Classification of sedimentary basins

Stretching the lithosphere: Rift basins

Nascent, conjugate passive margins 

Thrust faults: Some common terminology

Basins formed by lithospheric flexure

Basins formed by strike-slip tectonics

Allochthonous terranes: suspect and exotic

Source to sink: Sediment routing systems

Geofluids: Lithosphere-scale fluid flow

Geofluids: The permeability of faults

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Geofluids: Lithosphere-scale fluid flow

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Part of the Alberta Front Ranges fold -thrust belt north of Highwood Pass. Sawtooth-like Lower Paleozoic carbonates are exposed as flatirons in the hanging wall of a thrust immediately east of Lewis Thrust. Late Jurassic – Early Cretaceous foredeep deposits in the valley floor (mostly covered) were involved in the deformation as the orogenic load migrated eastward. Orogenic transport to the right.

The existence of fold-thrust belts in contractional orogens owes much to the presence of fluids and elevated pore pressures that reduce the critical stress necessary for faulting to occur. Part of the Alberta Front Ranges fold -thrust belt north of Highwood Pass.

Fluid flow at lithospheric scales

There’s not much that goes on in sedimentary basins that doesn’t involve fluids, particularly the aqueous kind. In sedimentary basins, subsurface fluid flow governs sediment compaction, the transfer of dissolved mass that produces a swath of diagenetic and metamorphic changes, heat flow, and mineralization. At the shallowest crustal levels, fluid flow is responsible for supplying groundwater to a third of Earth’s population.

Beyond the confines of sedimentary basins, fluids transfer mass (fresh water, carbon dioxide, methane, hydrocarbons, dissolved mass) and heat throughout the crust and lithosphere mantle where they play a significant role in dynamic, magmatic, and chemical processes. Water and its dissolved constituents, particularly carbon as aqueous CO2 and carbonate, are recycled from oceanic crust to the mantle lithosphere via subduction zones. Some of this interstitial water will be recycled during compaction, and some will be incorporated into the crystal lattices of common minerals such as clays, amphiboles and various micas, that in turn may be released during high temperature dehydration reactions, and eventually expelled to the atmosphere and hydrosphere.

Fluid flow of volatiles in the lower crust – upper mantle is slow, mostly accomplished along grain and crystal interfaces. Despite this seeming lethargy, these deep fluids are capable of transferring huge volumes of dissolved mass during metamorphic reactions.

The role of fluids and fluid flow in crustal and lithosphere-scale processes is illustrated using two examples: thrust faulting, and subduction zone fluids and magmas.

 

The mechanical paradox of thrust faulting

Dynamic processes like faulting are enhanced by the presence of fluids at elevated pressures. This is an important concept, first recognized by Hubbert and Rubey (1959 – PDF available). The starting point for their sophisticated analysis was the dilemma presented by the mechanics of overthrusting along shallow-dipping fault planes. Thrust faults can transport thick panels of rock many 10s of kilometres horizontally. The basic mechanical requirement for thrusting to occur is sufficient horizontal force to overcome frictional forces along the fault plane. And herein lies the dilemma – for thrusting to occur, the magnitude of these horizontal forces is so high that the rock body would disintegrate rather than being transported as a coherent structural unit. Hubert and Rubey’s task was to identify a mechanism that reduces friction to the point where real-world forces could be expected to do the job.

Our starting point is to look at fluid pressures at depth, for example in oil wells, represented by the expression:

                                                               P =  ρw gz

where P is the pressure of interstitial fluids at some depth measured vertically in the borehole, ρw is the density of the fluid (water), g = the gravitation constant, and z the depth from the surface to the point of interest (this expression is derived from Bernoulli’s equation for hydraulic potential). The expression is commonly referenced to two standard conditions:

  1. Hydrostatic pressure: the pressure at depth that supports a column of water extending to the surface (or close to the surface); in this case the pressure at any particular depth is defined by the weight of that water column where the average density ρ is close to one, and
  2. Lithostatic pressure (or geostatic pressure) that at any depth represents the weight of the overlying column of rock + water. In this case, density ρ is the average bulk density.
The concept of hydrostatic and lithostatic pressure at a depth ‘z’ represented as columns of water and rock respectively. Assuming the columns have unit area (area = 1) means that their volumes can be expressed as units of depth in the expression P = ρw gz.

The concept of hydrostatic and lithostatic pressure at a depth ‘z’ represented as columns of water and rock respectively. Assuming the columns have unit area (area = 1) means that their volumes can be expressed as units of depth in the expression P = ρw gz.

Hydrostatic fluid pressures tend to persist to 3-4 km depth, but at greater depths the pressures deviate from the hydrostatic trend – usually increasing (but in some situations can decrease). This condition is overpressured. In some cases, fluid pressures can approach or exceed lithostatic pressures. Fluid pressures exceeding hydrostatic values have been observed in many deep wells.

A typical pressure-depth curve for the Gulf Coast Basin shows the transition to elevated pore pressures at about 3000 m depth. Compaction disequilibrium plus significant changes in the diagenetic environment (e.g. quartz precipitation, hydrocarbon maturation) are responsible for changes in permeability and fluid transmissibility.

A typical pressure-depth curve for the Gulf Coast Basin shows the transition to elevated pore pressures at about 3000 m depth. Compaction disequilibrium plus significant changes in the diagenetic environment (e.g. quartz precipitation, hydrocarbon maturation) are responsible for changes in permeability and fluid transmissibility.

The expression P = ρgz can be rewritten in terms of the component of total vertical lithostatic stress Sz:

Sz = ρgz for a water-fluid column of unit area (Pressure = force per unit area, such that P = Sz/1, or P = Sz).

However, because Sz is the sum of the interstitial fluid pressure P plus the weight of the overlying solid rock column (Hubbert and Rubey call this the residual solid stress σz), then:

                                                                      Sz = P + σz

Hubbert and Rubey demonstrated that for overthrusting to occur, the critical shear stress τ (i.e., the minimum shear stress required to initiate movement across a plane) is equivalent to the vertical residual solid stress σz times a measure of material strength referred to as the angle of internal friction Tanθ:

                                                                  τ = σz Tanθ

(Tanθ is a rock or material property that refers to its ability to resist deformation and is measured as the angle between the normal stress and a resultant stress at the point where shear begins. The analogous measure in loose sediment or soil is the angle of repose, where slopes greater than this angle become unstable).

Thus, the critical stress at which overthrusting begins can now be written as:

                                                           τ = (Sz – P) Tanθ

This all-important equation indicates that as fluid pressure P increases and approaches the value of Sz, the critical shear stress τ approaches zero. Herein lies the elegance of Hubbert and Rubey’s analysis. As fluid pore pressures along the fault plane increase above hydrostatic values, friction is reduced to the point where overthrusting can occur with relative ease. Tectonic compression, differential compaction, and mineral dehydration reactions are some of the more common mechanisms that lead to increased fluid pressures.

The mechanical dilemma presented by overthrusting is resolved!

 

Fluid flow and partial melting in subduction zones

The production of magmas in the mantle is strongly dependent on the availability of water. Flux melting occurs when free water is available at solidus temperatures deep in the crust – upper mantle (the solidus is the temperature at which melting begins – it also defines the lithosphere-asthenosphere boundary). The water acts as a flux, lowering the melting point of different mineral components, thus promoting partial melting. For example, dry granite melts between 1100 – 1250oC, but in the presence of water melting can begin at temperatures as low as 650oC. This is illustrated in the graph below, where the solidus intersects the lithosphere geotherm at progressively lower temperatures, depending on the water content.

The mantle solidus curve (i.e., the temperature at which partial melting begins), moves towards the geotherm in concert with increasing water content. Partial melting begins when the two curves intersect.

The mantle solidus curve (i.e., the temperature at which partial melting begins), moves towards the geotherm in concert with increasing water content. Partial melting begins when the two curves intersect.

Flux melting is an important process in subduction zones. The water required to initiate melting is derived from the descending oceanic crust that contains wet sediment and wet basaltic volcanic and intrusive rock. As subduction proceeds, increasing compaction of the sediment matrix drives interstitial fluid flows to the upper plate, while increasing temperatures promote dehydration reactions in minerals that release water bonded to crystal lattices. Dehydration can begin at temperatures as low as 60oC in clays (Saffer and Tobin, 2011; PDF available); hydrocarbon maturation also begins at about this temperature and becomes increasingly rapid and pervasive at 80o-120oC (the oil generation window).  Fluid buoyancy plays a major role in its ascent. Thus, fluids in the subducting slab are partitioned between the upper plate or recycled in the underlying mantle.

The fate of these fluids is a topic of active research because they are implicated in the tectonics of subduction zones, particularly mega-earthquakes and slow-slip  along the subduction interface, the evolution of magmatism and volcanic arcs, and potential mineralization. Elevated pore fluid pressures resulting from tectonically induced compression also play a role in the generation of accretionary wedge thrust faults.

The origin and fate of fluids generated by subduction of oceanic lithosphere, the partial melting of asthenosphere mantle, and the formation of a magmatic arc. Some fluid in the oceanic crust is recycled to the deep mantle. Fluid flow in the accretionary prism is driven by compaction, tectonic compression, and clay dehydration. Note the deflection of geotherms by descending, cold, oceanic lithosphere. Modified from Farsang et al., 2021 and Miller 2013.

The origin and fate of fluids generated by subduction of oceanic lithosphere, the partial melting of asthenosphere mantle, and the formation of a magmatic arc. Some fluid in the oceanic crust is recycled to the deep mantle. Fluid flow in the accretionary prism is driven by compaction, tectonic compression, and clay dehydration. Note the deflection of geotherms by descending, cold, oceanic lithosphere. Modified from Farsang et al., 2021 open access; and Miller 2013.

Flux melting at the lithosphere-asthenosphere boundary creates partial melts that rise buoyantly through the upper plate; their ultimate fate is eruption at the surface and the construction of magmatic arcs. Huge volumes of fluid are also expelled to the atmosphere during eruptions – mostly water, plus lesser quantities of CO2, SO2, and inert gases such as helium. Distributed fluid expulsion at the sea floor also occurs via fracture networks and faults through the upper plate, and from compaction of the accretionary wedge. All these fluid expulsion processes contribute significant volumes of dissolved solids to the oceans.

A notable feature of volcanic arcs is their restriction to relatively narrow, linear regions of the upper plate. This implies some kind of mechanism that focuses both aqueous fluid flow and rising magmas; for example, fault and fracture networks, permeability channels, or the channelling of buoyant fluids via temperate gradients or changes in crustal rheology (e.g., Wilson at al. 2014; PDF available).

 

Other posts in this series

Sedimentary basins: Regions of prolonged subsidence

Defining the lithosphere

The rheology of the lithosphere

Isostasy: A lithospheric balancing act

The thermal structure of the lithosphere

Classification of sedimentary basins

Stretching the lithosphere: Rift basins

Nascent, conjugate passive margins 

Thrust faults: Some common terminology

Basins formed by lithospheric flexure

Basins formed by strike-slip tectonics

Allochthonous terranes: suspect and exotic

Source to sink: Sediment routing systems

Geofluids: Sedimentary basin-scale fluid flow

Geofluids: The permeability of faults

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