physical characteristics of sand injectites

Earth-Science Reviews 106 (2011) 215–246

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Physical characteristics of sand injectites

Andrew Hurst a,⁎, Anthony Scott a,b, Mario Vigorito a,c

a University of Aberdeen, Department of Geology & Petroleum Geology, King's College, Aberdeen, AB24 3UE, UKb Statoil, Reservoir Modeling, Sandslihaugen 30, 5254, Bergen, Norwayc Statoil, Exploration, Forus Vest, Stavanger, Norway

⁎ Corresponding author. Tel.: +44 1224273713; fax:E-mail address: [email protected] (A. Hurst).

0012-8252/$ – see front matter. Crown Copyright © 20doi:10.1016/j.earscirev.2011.02.004

a b s t r a c t
a r t i c l e i n f o
Article history:Received 21 December 2009Accepted 5 February 2011Available online 22 February 2011

Keywords:sand injectitesandstone intrusionssand injectionsand fluidizationsand extrusionoverpressurehydrofractureerosionflow processes

Almost two hundred years of research is reviewed that focuses on the physical characteristics of sandstoneintrusions. It is concerned with mechanisms of sand injection, particularly with fluid-grain transport andsedimentation processes during the remobilization, injection and extrusion of sand. Outcrop and subsurfacestudies in combination with laboratory experimental data are drawn on to present the state-of-the-art of sandinjection. The text covers 1) geometry, internal structure, and microtexture of deformed parent units, injectedand extruded sandstones, 2) host-strata and their seal characteristics that contribute to basin-wideoverpressure generation, 3) common trigger mechanisms for sand injection such as high magnitudeseismicity and the rapid injection of large volumes of fluids, 4) fluid types that drive sand into fractures, 5)hydrofracture mechanisms that induce regional-scale seal failure, 6) liquefaction and fluidization processesthat transport sand into fractures, 7) sedimentation processes in fractures, 8) the flow regime of fluidized sandduring injection, 9) post-sand-injection fluid flow and diagenesis, 10) porosity and permeabilitycharacteristics of injected sandstones and 11) post-sand-injection fluid-flow over geological timescales.Processes of sand remobilization, injection, and extrusion are complex and depend on many interrelatedfactors including: fluid(s) properties (e.g. pressure, volume, composition), parent unit and host-stratacharacteristics (e.g. depositional architecture, grain size and distribution, clay-size fraction, thickness,permeability) and burial depth at the time of injection. Many studies report erosional contacts between hoststrata and injected sands and these record high-velocity, erosive flow during injection. The flow regime ispoorly constrained and similar features are interpreted as records of laminar and turbulent flow, or both,during injection. Internal structures are common in sandstone intrusions and can be accounted for by a varietyof processes. The interpretational limits largely result from a lack of laboratory experiments that focus ondeveloping analogues for sand injection. The relationship between grain fabric developed during injectionand its control on permeability in sandstone intrusions is poorly understood and failure to advance this field ofresearch will hinder the quantitative characterization of sandstone intrusions as fluid-flow conduits duringbasin evolution. We conclude that future research should focus on: 1) quantification of sediment transportmodes under different flow conditions in different fracture dimensions with laboratory data relevant to sandinjection; 2) estimation of the effect of injection on the bulk permeability of otherwise low-permeability seals(host strata) so that their effect on fluid flow can be assessed at all scales; and 3) incorporation of sandinjection into quantitative basin models. Although an enormous amount of data have arisen from existingstudies there remains a need to advance many fields of research related to sand injection so that thesignificance of these important structures can be fully appreciated in the geological record.

Crown Copyright © 2011 Published by Elsevier B.V. All rights reserved.

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2162. Architectural elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 217

2.1. Sandstone dikes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2172.1.1. Injection geometry and margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2172.1.2. Internal sedimentary structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2182.1.3. Microtextures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 219

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2.2. Sandstone sills . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2212.2.1. Injectite geometry and margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2212.2.2. Internal sedimentary structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2212.2.3. Microtextures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 222

2.3. Columnar intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2232.3.1. Geometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2232.3.2. Internal structures and microtextures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 224

2.4. Irregular sandstone intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2242.4.1. Geometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2242.4.2. Internal structures and microtextures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 224

2.5. Sandstone extrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2242.5.1. Geometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2242.5.2. Internal sedimentary structures and microtextures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 224

2.6. Parent units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2272.6.1. Geometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2272.6.2. Internal sedimentary structures and microtextures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 228

2.7. Architecture of sand injectites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2292.8. Host strata: sealing capacity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 229

3. Sand injection: triggers, fracturation, and grain fluidisation and sedimentation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2303.1. Trigger mechanisms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2303.2. Fluid types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2313.3. Hydrofracturing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2313.4. Transport and sedimentation: liquefaction and fluidization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 232

3.4.1. Liquefaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2323.4.2. Fluidization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2323.4.3. Sedimentation during sand injection . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 233

3.5. Fluidized-flow regime during sand injection . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2333.6. Process of sand injection: overpressure development, liquefaction, hydrofracturing and fluidization . . . . . . . . . . . . . . . . . 234

4. Post-sand-injection fluid flow and diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2355. Petrophysical characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2356. Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 235

6.1. Overview . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2356.2. Identification of sand injectites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2376.3. Injectite margins and internal structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2376.4. Sand extrusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2386.5. Implication of sand injectites on post-sand-injection fluid-flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 239

7. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 239Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 239Appendices: synthesis of literature review . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 240References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 243

1. Introduction

Study of sand injectites is relatively new in the geological literaturealthough the first description of a sandstone dike dates back almost200 years Murchison (1827). Since that early description sandstonedikes were an irregular subject of papers (e.g. Diller, 1890; Newsome,1903; Jenkins, 1930; Smyers and Peterson, 1971) until their significancein hydrocarbon reservoirs was recognized by Dixon et al. (1995). Inparticular from the interpretation of seismic reflection data used forhydrocarbon exploration and production (MacLeod et al., 1999; Lone-rgan and Cartwright, 1999) other sandstone geometries than relativelyplanar dikes were identified including conical intrusions (Huuse et al.,2005) and saucer-shaped intrusions (Hurst et al., 2003a), someofwhichconstituted sufficiently large volumes of reservoir-quality rock to betargets for drilling in oilfieldsunderdevelopment (MacLeod et al., 1999;Duranti et al., 2002a), including deliberate exploration targets (Raw-linson et al., 2005; De Boer et al., 2007; Szarawarska et al., 2010). Theknowledge of sandstone intrusions from the subsurface is arguablybetter than from outcrops.

Sand injectites have several components not all of which aresandstone intrusions that comprise sills, high- and low-angle dikes andoccasionally features with irregular geometry (Thompson et al., 1999;Thompson et al., 2007; Scott et al., 2009) and rarely diapiric features(Williams, 2001), and depositional sandstones with varied levels ofdeformation and referred to as parent units (Vigorito et al., 2008),Additional components of sand injectites recognized since Hurst et al.

(2003a) include sandstone extrusions (termed extrudites, Hurst et al.,2006), and hydrofractured mudstones (Vigorito et al., 2008; Vigoritoand Hurst, 2010). The components of sand injectites are geneticallyrelated and dikes and sills combine to form conical and saucer-shapedintrusions kilometer-scale in lateral extent, 10's of meters thick andoften cross cut N100 m of stratigraphy ( Huuse et al. 2005; Huuse et al.2007; Vigorito et al., 2008; Vigorito and Hurst, 2010) as well as formingsmall individual features (b0.01 aperture and b0.1 length).

Sand injectites are recognized in many geodynamic and geologicalsettings globally, throughout the stratigraphic record (papers in Hurstand Cartwright, 2007), and typically are fine- to medium-grainedsiliciclastic sand although much coarser-grained (Hubbard et al., 2007)and carbonate sand examples (Hurst and Cartwright, 2007) occur. Sandinjection occurs when fluidized sand is forced into host strata and fills anetwork of hydrofractures thus creating a network of intrusions (Hurstet al., 2003a). There is however a limited understanding of the flowprocesses during injection, themost robust direct evidence coming fromobservations of structures that formed along the margins betweensandstone intrusions (e.g. dikes, sills) and host strata (Diggs, 2007; Scottet al., 2009). Sandstone intrusions are most easily identified by theirrelationshipswith host strata when they formdiscordant or concordantelements that branch, split, and rejoin (Archer, 1984) or form otherbedding-discordant geometries (Diggs, 2007).

Here we review published data on the physical character ofremobilized and injected sand. In particular the review will integrateoutcrop andmicrotextural data, and then compare thesewith theoretical

217A. Hurst et al. / Earth-Science Reviews 106 (2011) 215–246

and experimental considerations of sand fluidization in order to describeand discuss the following topics: 1) characterization of externalgeometries and internal structure of remobilized and injected sandstone;2) grain-scale processes that affect the (petro) physical characteristics ofsandstone intrusions; and3) sandfluidizationprocesses. This reviewalsoprovides a comprehensive database (Appendices A–E) whereby readerscan access studies that focus on the geometry, internal structures,petrography, and physical character of sandstone intrusions. It is beyondthe scope of this paper to provide a detailed review of subsurface sandinjectites; documentation of the interpretation of seismic reflection dataHuuse et al., 2007) and some aspects of borehole interpretation (Durantiand Hurst, 2004) are found elsewhere. Readers interested in therelevance of sand injectites to hydrocarbon exploration and production,are referred to Hurst and Cartwright (2007a).

2. Architectural elements

Sandstone intrusions and remobilized sandstones exhibit a range ofgeometries but can be broadly subdivided into four elements: parentunits, dikes, sills and extrudites. Parent units are depositionalsandstones (Fig. 1) that display features formed both by depositionalprocesses and post-depositional sand and fluid mobilization, and forman interconnected system of sandstones together with sandstoneintrusions. Sandstone dikes (red arrows; Fig. 1) are discordant,sometimes tabular bodies that may include mudstone clasts, organic

Fig. 1. A synoptic of a sand injectite complex with a tripartite architecture based on outcrop ahost rock (dark grey). Remobilized parent sandstone units (yellow arrows); sandstone dikesandstone extrudites (green arrow).

matter and diagenetic cements, Truswell, 1972) that over the majorityof their length clearly crosscut sedimentary bedding at low and highangles, arbitrarily defined as b20º and N20º), respectively. Sandstonesills (blue arrows; Fig. 1) are crudely tabular bodies of injected materialwhose boundaries are approximately concordant with that of beddingin the host strata but at some points are discordant with bedding alongtheir upper and lowermargins Vigorito et al., 2008). Irregular sandstoneintrusions (orange arrow; Fig. 1) have highly discordant margins andrapid changes in thickness. Extrusive sandstones (extrudites, greenarrow; Fig. 1) formed by venting of sand on paleo-surfaces such as theseafloor and were connected to and fed by underlying dikes.

2.1. Sandstone dikes

Sandstone dikes are recognized by their discordance with andtruncation of bedding in host strata and they contain a variety ofinternal sedimentary structures (Appendix A).

2.1.1. Injection geometry and marginsTypically, sandstone dikes crosscut host-strata bedding at high-

angles (Fig. 2A) but low angle dikes are common and often large(N10 m) aperture (Fig. 2B). Dikes display a range of geometries(Appendix A) from straight and planar (Taylor, 1982; Fig. 3A) tohighly irregular, bulbous and curved (Parize et al., 2007a; Fig. 3B);individual dikes may include several geometric styles (Surlyk et al.,

nd subsurface observations (substantially modified after Hurst and Cartwright, 2007) ins (red arrows) and sills (blue arrows); irregular sandstone intrusions (orange arrow);

image of Fig.�1

Fig. 2. Geometries of sandstone dikes; lithologies represented are mudstone (dark grey) and sandstone (light grey). A) Steep high-angle dikes from the upper part of the intrusivecomplex in Marca Canyon, Panoche Hills (see Hurst et al., 2007, Fig. 15; Vigorito et al., 2008). B) A large (8–12 m aperture) low-angle dike from Right Angle Canyon (Vigorito et al.,2008; Vetel and Cartwright, 2009, Fig. 6). C) Vertical to sub-vertical dikes with straight and planar margins. These dikes display horizontal offsets at several levels along beddingplanes in the host mudstones (Alexander Island, Antarctica, redrawn andmodified from Taylor, 1982). D) Ptygmatic folding of a dike with numerous host mudstone clasts. Folding ofthe dike is attributed by Parize et al. (2007a) to post-sand-injection compaction during burial (Vocontian Basin, France, redrawn and modified from Parize et al., 2007a); sandstonedikes (light grey) in host mudstones (dark grey).

1 For convenience all clasts of fine-grained material are termed mudstone clasts.They are commonly eroded from host strata as mud or very poorly consolidatedmudstone and subsequently have consolidated during burial.


2007). Other common geometries include tapering (Rodriguez-Pascua et al., 2000; Strachan, 2002), bifurcating (Surlyk and Noe-Nygaard, 2001; Hubbard et al., 2007) and ptygmatic features,including folds (Surlyk and Noe-Nygaard, 2001; Satur and Hurst,2007) that may be caused by post-sand-injection compaction of thehost strata (Hillier and Cosgrove, 2002).

The range of sandstone dike geometry, from planar to curvilinear, isdescribed by many authors (e.g. Taylor, 1982; Surlyk and Noe-Nygaard,2001). Planar dike geometry is clearly suggestive of sand injection alongplanar fractures.However, the formationof irregular, bulbous, and curveddikes is more difficult to explain in terms of sand injection into fracturesystems. Truswell (1972) attributed irregularities in dike geometry toanisotropy in the host strata and post-injection folding and boudinage.

Discordant marks are common on the margins of sandstone dikesand include flutes or flute-like marks (Peterson, 1968; Hillier andCosgrove, 2002; Hubbard et al., 2007), grooves (Taylor, 1982; Surlykand Noe-Nygaard, 2001), rills (Diggs, 2007), lobate scours (Lewis,1973; Martill and Hudson, 1989), frondescentmarks (Surlyk and Noe-Nygaard, 2001) and gutter marks (Keighley and Pickerill, 1994). Therelief of the marks ranges from millimeter to centimeter scale. Thepresence of discordant marks along dike margins provides a record ofhow pressurized, fluidized sand interacted with the surfaces offractured host strata during injection. From papers above it is unclearwhether the formative processes of these structures are similar tothose of sole marks in turbidites or whether other mechanisms suchas frictional drag along the dike margins play a role. Keighley andPickerill (1994) described similar curved and bulbous marks (termed“gutter-like” marks) on the margins of clastic dikes, which theyattributed to rheoplastic molding of the host strata; a similarconclusion was reached by Surlyk et al. (2007).

2.1.2. Internal sedimentary structuresInternal sedimentary structures are a distinct and pervasive

feature of sandstone dikes (Fig. 4; Appendix A). Most commonly,

they display banding that ranges in thickness from 0.01 to N0.10 m(Diggs, 2007; Fig. 4A), which is usually bounded by sharp contactsthat are defined by differences in grain size (Archer, 1984), grainalignment (Hillier and Cosgrove, 2002), and clay-size particles (Diggs,2007; Fig. 3A). Bands are commonly oriented parallel to the dikemargins (Gożdzik and Van Loom, 2007) but also form perpendicularto dike margins (Truswell, 1972; Diggs, 2007; Fig. 3C).

Laminae are recorded (Hubbard et al., 2007; Macdonald andFlecker, 2007), typically parallel to dike margins and, in commonwithbanding are defined by sharp boundaries created by differences ingrain size (Taylor, 1982), grain alignment, and concentrations of clay-size particles (Diggs, 2007; Fig. 4C). Individual laminae are typicallyb1 mm thick a characteristic that differentiates them bands; thesimilarity in the physical character of both bands and laminae issuggestive of a similar origin. Formation of banding and laminae areattributed to multiple episodes of injection (Taylor, 1982; Diggs,2007), irregularities in the rate of opening of fractures (Peterson,1968), or variations in both the velocity and viscosity of the injectingsand-charged fluids (Peterson, 1968; Taylor, 1982; Diggs, 2007).

Grading is common in sandstone dikes and occurs in two distinctstyles: approximately perpendicular to dike margins and along thelength of dikes; normal (i.e. the coarsest grains are found adjacent tothe dike walls) and reverse grading occur (Hubbard et al., 2007).Grading along the length of sandstone dikes occurs where there tendsto be an overall fining-upward (Obermeier, 1996; Ross and White,2005; Hubbard et al., 2007). Grain-size variation occurs as a functionof sand composition in parent units but also because of the presence ofclasts derived from fractured and eroded host strata. The latter arepredominantly mudstone1 clasts althoughwhere sandstone is presentin the host-strata sandstone clasts may occur (Glennie and Hurst,

image of Fig.�2

Fig. 3. Internal sedimentary structures in sandstone dikes in a mudstone host. The dikes are locally enriched with clay-size particles, in otherwise clean sandstone (light grey).A) Stratified sandstone dike characterized by light and dark bands the latter that correlate with concentrations of detrital clay-size particles. Other internal sedimentary structuresinclude stress pillars and flame structures (subvertical white structures) (Marathon Basin, Texas, redrawn andmodified from Diggs, 2007). B) Internal sedimentary structures foundnear the margins of a sandstone dike that include abundant mudstone clasts and sandstone enriched in detrital clay-size particles. The clay particles were produced by erosionalscouring of the host mudstone during sand injection (Marathon Basin, Texas, redrawn and modified from Diggs, 2007). C) A slightly oblique section through a sandstone dike thatcuts bedding at approximately 80°. Banding (0.1 to 0.25 m thick) is the dominant internal structure. The right-hand dike margin has a low-amplitude wavy surface with abundanterosive markings including striations and a polished surface; the left-hand side of the dike is more irregular and less well-exposed. From the Panoche Giant Injectite Complex,Moreno Formation (Vigorito et al., 2008); hammer length 0.4 m.


2007). Most examples of sandstone dikes consist of fine- to medium-grained sand and this grade of sand is known to fluidize preferentially(Lowe, 1975). This narrow range in grain size discourages grading butwhen a range in grain size occurs, for example, the clastic dikes of theMagallenes Basin (Hubbard et al., 2007), spectacular grading occurs inwhich conglomerate grades into medium-grained sandstone. Gradingof mudstone clasts perpendicular to dike margins is common and isattributed to sorting of eroded clasts of host strata during sandinjection (Obermeier, 1996; Diggs, 2007).

The processes responsible for grading in sandstone dikes arepoorly constrained but when assessing grading across a dike it is likelythat these processes are different to those responsible for gradingalong a dike. Taylor (1982) suggests that in the central parts ofsandstone dikes the segregation of coarser-grained mudstone clastsderived from the erosion of host strata may be indicative of laminarflow during sand injection. Experiments demonstrate that in laminarflow the larger solid-particles separate from dike walls and tend tosegregate toward the centre of an injected sheet (Bhattacharji andSmith, 1964). Hubbard et al. (2007) conclude that the normallygraded character of sandstone dikes is evidence of turbulent flowconditions during sand injection.

2.1.3. MicrotexturesUntil recently little was published on microtextures in sandstone

dikes (Peterson, 1968; Truswell, 1972; Taylor, 1982; Thompson et al.,1999). Platy and elongate minerals (i.e. mica, elongate quartz grains,plagioclase and mudstone clasts) tend to be aligned parallel to dikemargins (Fig. 5), which Taylor (1982) suggests is a microtextureindicative of the upward injection of sand. However, grain alignmentis not present in all sandstone dikes and in some instances elongategrains are unaligned (Truswell, 1972).

It is important to recognize that there is currently no unequivocalevidence that grain alignment in sandstone dikes is indicative of aparticular flow regime and the factors that help develop (or inhibit)grain alignment during sand injection are poorly understood. It wasshown that the degree of grain alignment correlates inversely withthe amount of clay-size particles present. Sandstone dikes thatcontain high concentrations of clay-size particles (up to 42%, Diggs,2007) have little evidence of grain alignment. Dott (1966) inferredthat grain alignment was indicative of viscous-laminar flow duringsand fluidization and injection .This inference seems unwarranted inview of the fact that similar features are produced by turbulent flow indepositional settings (Lowe, 1975).

image of Fig.�3

Fig. 4. A vertically-oriented thin-section photomicrograph of vertical zoning found atthemargin of a sandstone dike. Three distinct zones coarsen inwards (toward the right)from the margin. A coarse sand-size zone of grains on the right is separated (dashed redline) by a fine sand-size zone of grains (modified from Taylor, 1982).

Fig. 5. A sill-dominated intrusive complex (Fort Genois, Miocene, Tabarka, Tunisia). (A) TheThe entire section is ~120 m thick and 30–50% of the section is sills. (B) Line-drawing trace othick) is a depositional sandstone with faint bedding surfaces preserved (light grey). Withprimary sedimentary structures are no longer identifiable (for geological setting see Parizebodies that are 0.1–2 m thick (dark grey). The sill zone is approximately ~100 m thick.


When consideringflowprocesses during sand injection arguably thesingle most important microtextures in sandstone dikes are associatedwith clasts of host strata (Fig. 3B, C, 4B). Individual clasts, which aretypically mudstone, vary from angular to rounded, very fine-sand-sizeto boulder-size, and can be elongate or equant. The clasts are not usuallydistributed evenly within sandstone dikes and clasts are often mostabundant toward the margins of dikes (Diggs, 2007; Hubbard et al.,2007;Macdonald and Flecker, 2007; Scott et al., 2009) or at the junctionbetween dikes and sills (Parize et al., 2007a). In a few examples theproportion of clasts of host-strata increases toward the centre of a dike(Dixon et al., 1995). Vertical gradation of host-strata clasts may occuralong the long axis of sandstone dikes (Obermeier, 1996).

Although it is generally believed that mudstone clasts are derivedfrom the host strata several mechanisms for their formation areconsidered possible. Rafts of host strata are stoped from the dikemargins and incorporated into the fluidized sand during injection andprogressively disintegrate thereby introducing finer particles (Diggs,2007). Diggs does not discuss the precise processes responsible for thedisintegration of the rafts but it may be a combination of micro-fracturingorerosionby the abrasiveflowoffluidized sand.Other studiessuggest that clasts are derived by rip-up (Hubbard et al., 2007) or rip-down (Chough and Chun, 1988; Hamberg et al., 2007) from strata alongthe dike margins. Clasts of host strata in dikes are generally consideredto be locally derivedparticularlywhen associatedwith “jigsaw” textures(Duranti and Hurst, 2004) that are unlikely to be preserved unlesslocated within a meter or less from their point of origin. At least oneinstance is knownwhere, byusingbiostratigraphic data, it can be shownthat aminor proportionofmudstone clastswithin amudstone-clast richunit were derived and transported upward ~200 m (P.B. Rawlinson,personal communication, 2006).

In sandstone dikes in which clasts of host-strata are common thepores are usually enriched with clay-size particles (Truswell, 1972;Duranti and Hurst, 2004; Diggs, 2007). Pore-filling clay-size particlescan locally form up to 42% of thewhole rock volume (Diggs, 2007) andare most abundant close to the margins of sandstone dikes (Diggs,2007) forming a matrix that partly obstructs pore throats (Durantiand Hurst, 2004). It is important to recognize that, although most

section youngs to the left and is dominated by sills with minor dikes connecting them.f the complex. The lowermost sandstone shown in the right-hand field of view (N15 min the sandstone the degree of modification by fluidization increases upward where, 1988). The sill zone (dashed black line) is dominated by bed-concordant sandstone

image of Fig.�4

image of Fig.�5


pore-filling clays are detrital, some clay-size particles may beauthigenic (Jonk et al., 2003). The distinction between authigenicand detrital pore-filling clay-size particles becomes important whenassessing the influence of clay-sized material on flow processes. It iswell known that clay-size particles create significant intergranularcohesive forces (Middleton and Hampton, 1976; Amy et al., 2006) thatmay increase the minimum-fluidization velocity of sediment (Lowe,1975; Diggs, 2007), and can cause granular material to behavethixotropically (Truswell, 1972).

2.2. Sandstone sills

Sandstone sills display localized discordance with bedding,without which they are similar in appearance to some depositionalsandstones, particularly when the latter are modified by sandfluidization. Differentiation of sills from depositional sandstones byobservation of internal structures alone, for example in cores, is oftenchallenging and because of this we believe that many sills are“missed” in subsurface analysis (Appendix B).

2.2.1. Injectite geometry and marginsSandstone sills form elements of sand injectites that are approxi-

mately concordant with host strata (Fig. 6) but they may bifurcate(Truswell, 1972; Hillier and Cosgrove, 2002; Diggs, 2007; Lonerganet al., 2007; Parize et al., 2007a), taper (Hillier and Cosgrove, 2002;Kawakami and Kawamura, 2002; Lonergan et al., 2007; Parize et al.,2007a) and step (Fig. 7D;Hiscott, 1979; Archer, 1984; Obermeier, 1996;Huuse et al., 2004; Macdonald and Flecker, 2007). Three distinct sillgeometries are recognized: staggered, stepped and multi-layered thatoccur in this order from deeper to shallower in the ~350 km2 exposureof the Panoche Giant Injectite Complex (see Section 2.6; Vigorito et al.,2008; Vigorito and Hurst, 2010). Because of the limited lateral andvertical exposure of other sand injectites similar organization of sills isnot recognized elsewhere although the individual characteristics andgeometries are. Individual sills described elsewhere range fromb0.001 m (Surlyk and Noe-Nygaard, 2001) to 0.5 m (Hillier and

Fig. 6. Geometries of sandstone sills; lithologies represented are mudstone (dark grey) and saand split along bedding planes in the host mudstone (Quebec Appalachians, Canada, redrawand irregular geometries with rounded margins (Jameson Land, East Greenland, redrawn andisplays extreme pinch-and-swell thickness variations (Jameson Land, East Greenland, redrplanar margins (dashed black lines) that step up in the host mudstone (Tumey Hills, Eocen

Cosgrove, 2002), N3 m (Hiscott, 1979), and N12 m (Parize et al.,2007b); their lateral extent may exceed ~100 m (Parize et al., 2007a)but typically they extend b25 m (Obermeier, 1996). Rapid lateralchanges in thickness (Surlyk andNoe-Nygaard, 2001) or extremepinch-and-swell features (Hiscott, 1979; Diggs, 2007) are common. Planar(Truswell, 1972; Hiscott, 1979; Parize et al., 2007a; Fig. 6A) curved(Surlyk and Noe-Nygaard, 2001; Fig. 6B, C) and, irregular (Surlyk et al.,2007; Fig. 6B) margins occur that are often erosive.

Erosional features, specifically irregular margins Fig. 6) are commonalong themargins of sills (Obermeier, 1996; Kawakami and Kawamura,2002; Diggs, 2007; Macdonald and Flecker, 2007; Vigorito et al., 2008;Scott et al., 2009; Kane, 2010) and occur on both the upper (Archer,1984; Kawakami andKawamura, 2002; Diggs, 2007) and lowermargins(Kawakami and Kawamura, 2002; Macdonald and Flecker, 2007).Obermeier (1996) reported upward erosion 0.2 m into the overlyinghost strata whereas Kawakami and Kawamura (2002) reporteddownward erosion of 0.05 m. Where large upward erosive featuresare recorded they are termed scallops (Hurst et al., 2005), arecharacteristic of stepped sills (Fig. 6D) and commonly erode severalmeters into mudstone overburden (4 m, Scott et al., 2009; 2 m Vigoritoet al., 2008). Estimates of erosion are typically minimum values as thelocation from which erosion started is rarely preserved.

2.2.2. Internal sedimentary structuresSandstone sills are variously reported as structureless (Truswell,

1972; Hillier and Cosgrove, 2002; Parize et al., 2007b) or as containinginternal structures (Kawakami and Kawamura, 2002; Macdonald andFlecker, 2007; Scott et al., 2009). Where present sedimentarystructures form by scouring (Fig. 7A), splitting (Fig. 7B, C), folding(Fig. 7B) and imbrication (Fig. 7C,D) of clasts of host (mudstone)strata (Kawakami and Kawamura, 2002); typically the eroded clastsare disintegrated and reworked into muddy laminae (Fig. 7C). In allcases there is an increase in clast content toward the sill margins(Fig. 6A, B, C) and the injecting sands have ripped and reworked thehost strata into clasts, forming normally-graded intervals (Macdonaldand Flecker, 2007).

ndstone (light grey). A) Sandstone sills with sharp planar margins that step up or downn and modified from Hiscott, 1979). B) Sills displaying both straight and planar marginsd modified from Surlyk et al., 2007). C) Irregular and rounded sandstone injectite thatawn and modified from Surlyk et al., 2007). D) Sandstone sill with irregular and sharpe, California, Huuse et al., 2004).

image of Fig.�6

Fig. 7. Internal sedimentary structures in sandstone sills; lithologies represented are mudstone (dark grey) and sandstone (light grey). A) Scoured erosional upper margin withnumerous mudstone clasts found adjacent to the margin (Schmidt peninsula, Sakhalin, redrawn and modified from Macdonald and Flecker, 2007). B) Elongate mudstone clastsaligned parallel to the lower host-mudstone margin (Jameson Land, East Greenland, redrawn and modified from Surlyk et al., 2007). C) Complex internal sedimentary structuresfound near a sill upper margin, which includes splitting and intrusion along host-mudstone bedding. Other internal structures include imbricated host-mudstone clasts, muddylaminae, and sandstone-filled microfractures in the mudstone clasts (Honshu Island, northeast Japan, redrawn and modified from Kawakami and Kawamura, 2002). D) Part of amultilayer sill in which some original mudstone bedding is preserved near the base but becomes increasingly disrupted and fragmented upward. Mudstone rafts are elongate,extending 1–2 m in length, and are cross-cut by sandstone-filled fractures, which fragment the rafts into smaller clasts (Moreno Gulch, Panoche Hills, California; Vigorito et al., 2008;Vigorito and Hurst, 2010).


The internal sedimentary structures (Fig. 7) are suggestive ofinternal flow and sedimentation processes, such as fluidized sand flowwithin a fracture. The presence of muddy laminae (Fig. 7C) andimbricated mudstone-clasts (Fig. 7C) also suggests that there weresustained periods of traction during the formation of sills (Kawakamiand Kawamura, 2002). Alternatively, they may have formed byshearing of concentrated liquefied sand that deforms by hydroplastic

Fig. 8. Thin-section photomicrographs of textures found in mudstone clasts within a sandsmudstone clasts (Md) that are pervasively cut by sandstone-filled microfractures. Some of th(black arrows). Both micrographs are modified versions of figures from Kawakami and Kaw

laminar flow (Lowe, 1976; Allen, 1984). This could have a variety ofcauses, including high sedimentation rates (Lowe, 1982) and a suddenloss of fluid pressure that froze the flow.

2.2.3. MicrotexturesPetrographic analysis reveals distinct microtextures in which

abundant clay-size particles and platy grains (i.e. elongate mudstone

tone sill. A) Host mudstone clasts (Md) cut by sandstone-filled microfractures. B) Hoste microfractures are filled by sand-size grains and smaller silt-size host mudstone clastsamura (2002), Honshu Island, northeast Japan.

image of Fig.�7

image of Fig.�8


clasts and mica) are aligned parallel to the margins of sills (Kawakamiand Kawamura, 2002; Diggs, 2007). Similar microtextures are reportedbyKawakami andKawamura (2002)with additionalmicrotextures thatinclude mudstone clasts, crosscut by sandstone-filled microfractures(white arrows; Fig. 9A, B; (Kawakami and Kawamura, 2002; Scott et al.,2009). Numerous mudstone clasts derived from host strata arecommonplace (Md; Fig. 8) that are interpreted to be evidence of erosionof the host strata during sand injection.

2.3. Columnar intrusions

Spectacular columnar sandstone intrusions are recorded in aeoliansandstones and smaller, similar intrusions are present in other

Fig. 9. A range of geometries in some columnar sandstone intrusions. A) Free-standing coluBasin, Middle Jurassic, SE Utah, modified from Huuse et al., 2005). B) Cliff face exposure wfaulted host strata of inter-bedded Entrada sandstone and siltstone (Shephard's Point, Kodacshowing a columnar sandstone and its relation to the host strata. Internally the pipe is apparsandstone (Navajo Sandstone, Jurassic, Utah, modified from Chan et al., 2007). The overall mof a columnar sandstone encased within host mudstone. The columnar sandstone has a spthickness decreases downward. Filamentous (b0.001 m thick) sandstones layers emanate f

depositional environments. Current published records of columnarsandstone intrusions are largely restricted to eolian strata (Appendix C).

2.3.1. GeometryClastic pipes and megapipes, here collectively referred to as

columnar intrusions, are recognized in the eolian Carmel Formationon the Colorado Plateau (Chan et al., 2007). They have columnargeometry with sharp margins that cut bedded strata at a high angles(Fig. 9A,B) and have flared bases.Where bedding is present in the hoststrata it dips away from the margins of intrusions (Davidson, 1967;Fig. 9B). Upward-flaring of pipes (Chan et al., 2007) is observed andappears very similar to the concave-upward vent fills below sandextrudites (Hurst et al., 2006; Pringle et al., 2007, Figs. 4–9) although

mnar sandstone with remnant host strata in the right-hand field of view (Kodachromeith a columnar sandstone (dashed black outline) and its relation to down-sagged andhrome Basin, Middle Jurassic, SE Utah, modified from Huuse et al., 2005). C) A cliff faceently structureless. Note that the pipe connects to an overlying unit of heavily-fluidizedorphology appears to be twisted with a spiral axial geometry. D) Core-scale observationiral axial geometry giving a cork-screw appearance similar but smaller than C. and itsrom the margins of the intrusion along partings in the mudstone.

image of Fig.�9


evidence of surface extrusion is not recorded in the former. Althoughmany of these intrusions are near vertical and columnar, lower-angleand more sinuous intrusions occur (Fig. 9C). Margins of intrusions arecommonly brecciated and rich in granule-sized clasts of host strata.Megapipes differ from pipes (sensu Chan et al., 2007) only because oftheir size and a tendency to contain extra-formational clasts. In theCarmel Formation columnar intrusions vary from b1 m to as much as100 m in diameter.

Columnar intrusions are recorded in sand injectites that thatintruded deepwater clastic mudstones (Fig. 10D) but we are unawareof columnar intrusions of similar size to those recorded from theCarmel Formation. We suspect that subsurface examples of largecolumnar-intrusions (megapipes of Chan et al., 2007) occur (e.g.Løseth et al., 2003) but remain unvalidated by borehole data. Thesmall curved columnar intrusion with external ornament (Fig. 10C)formed by erosion of the host mudstone in clockwise vortex flow thuscorresponding to positive vorticity (the mirror image of the anti-clockwise vorticity in the northern hemisphere atmosphere).

2.3.2. Internal structures and microtexturesMany columnar intrusions are reported as structureless (Chan et

al., 2007) but exposures of cylindrical margins do not favor viewing ofinternal structures. However, irregular and deformed subhorizontalbedding and lamination do occur and appear similar in scale (5–15 cmwidth) to banding observed in other sandstone intrusions (e.g. Scottet al., 2009). In plan-view nested structures formed by concentriccylindrical rings are observed and in large columnar intrusions(megapipes) may occur as concentric zones of angular conglomerate(breccias) that surround finer-grained sandstone; vertical trends ingrain size also occur. The presence of pervasive ring fractures or faultswas noted by Netoff (2002).

Microtextural data indicate that the average grain size in columnarintrusions is greater than in the strata they cut (Chan et al., 2007),which indicate that the intrusions are derived from the fluidization ofunderlying rather than laterally-adjacent strata. Compositionally thecolumnar intrusions are very similar to their host strata (Netoff andShroba, 2001) although Chan et al. referring to a pers comm.(Mahaney, 2004) record the presence of possible percussion marksas possible indicators of turbulent flow.

2.4. Irregular sandstone intrusions

Some sandstone intrusions have irregular external geometry andare clearly neither dikes nor sills. Mounded geometry interpreted aspart of sand injectites is commonly reported from the subsurface andknown from outcrops (Appendix D).

2.4.1. GeometrySandstone intrusions with irregular geometry are characterized by

discordant margins and rapid changes in thickness (Fig. 10). Irregularsandstone intrusions are characterized by discordant, erosive lower(Fig. 10A,B,F) and upper margins, the latter typically form convex-up“scallops” that often cut N1 m into overlying host strata (Fig. 10A, C)and are highly discordant with respect to bedding in the surroundinghost mudstones (Hurst et al, 2003a; Surlyk et al., 2007; Scott et al.,2009).

Convex-up upper margins may record sand diapirism into theoverlying mudstones (Williams, 2001; Surlyk et al., 2007) however, aproblem with this interpretation is that if only minor deformation isreported in the encasing host mudstones diapirism is an unlikelyprocess. Sand diapirism on this scale should produce extensivedraping and deformation of the surrounding host strata (cf. Williams,2001, Fig. 5f). Observations by Hurst et al. (2003a) support an erosiveorigin for the convex-upward mounds similar to those describedearlier from sills (scallops of Hurst et al., 2005) that in one examplemay have represented a 20–30 m high feature (Duranti et al., 2002b;

Hurst et al., 2007); a similar erosive origin is supported by Diggs(2007). An alternative interpretation is that discordant upper marginsform by post-sand-injection loading by the fine-grained overburdenbut it is highly unlikely that loading can cause an originallyconcordant structure to develop discordance along only its uppermargin (Diggs, 2007). Larger-scale mounded sandstone intrusions inthe subsurface offshore Denmark have similar erosional features(Hamberg et al., 2006, 2007).

2.4.2. Internal structures and microtexturesIrregular sandstone intrusionsmay be devoid of internal structures

(Surlyk and Noe-Nygaard, 2001; Lonergan et al., 2007) or containdiverse internal structures (Diggs, 2007; Hamberg et al., 2007; Scott etal., 2009). Most of the internal structures reported are defined byconcentrations of mudstone clasts, which lends support to the ideathat the irregular geometries are formed by the erosion of host strata;this hypothesis is independently supported by microtextural evi-dence. At granular scale the irregular sandstone bodies are charac-terized by poor sorting and high detrital-clay content (Diggs, 2007).Similar textures are recorded from mounded sandstone intrusions inthe subsurface (Hamberg et al., 2007) where numerous mudstoneclasts are observed near the margins. Individual mudstone clasts(N0.01 m diameter) contain smooth millimeter-scale incisions filledby sand grains. The incisions are interpreted as recording the erosiveinteraction between fluidized sand and the mudstone clasts duringsand injection (Hamberg et al., 2007; Scott et al., 2009).

2.5. Sandstone extrusions

Sandstone extrusions (extrudites, Hurst et al., 2006), are knownfrom a variety of stratigraphic settings and ages and in their simplestform occur as sand(stone) volcanoes or as composite, thicker, sheet-sand(stone) units (Appendix E).

2.5.1. GeometrySandstone extrusions are most commonly associated with conical

and sheet-like geometries (Fig. 11). Sandstone volcanoes are thesimplest geometry and typically have a conical to elliptical geometryand are sourced by individual dikes (Fig. 11A,B). Their knownthickness and diameter when observed at outcrop range from 0.1 to0.75 m and 0.3 to 3 m, respectively. Recently a possible “giant”sandstone volcano (80 m thick, 5.3×107 m3) was identified from thesubsurface (Andresen et al., 2009). Extruded sandstone sheets form adistinct geometry that may exceed 3 m thickness and extend laterallyfor several hundred meters (Boehm and Moore, 2002; Hurst et al.,2006) or even regionally (Vigorito et al., 2008).

Contact relationships between the sandstone extrusions and theunderlying substrate are often highly irregular and discordant(Fig. 11C) or bedding parallel (Fig. 11D). Discordance is especiallycommon where dikes reach the paleo-surface and the underlyingsubstrate may be highly disrupted by fractures and brecciation (Hurstet al., 2006). Down-warping of the lower contact toward ventingpoints is observed (Obermeier, 1996; Pringle et al., 2007) that appearto be sandstone fills of hollows excavated during vent formation andmay be related to pock marks (Andresen et al., 2009). More distalfrom venting points the contact relationships are sharp and nonerosive. In most cases the upper contact of sandstone extrusions isdraped by fine-grained strata and in general an overall fining-upwardis observed.

2.5.2. Internal sedimentary structures and microtexturesLow-angle bedding and laminae are common in sandstone extru-

sions (Appendix D). Typically the low-angle beds and laminae dip awayfrom a central venting point forming low-relief cones (Pringle et al.,2007; Jonk et al., 2007; Fig. 11D) however, when the sandstones arethick they often comprise interbedded and interfingering beds that

Fig. 10. A range of geometries present in some irregular sandstone intrusions. A) Yellowbank Creek sand injectite (Miocene) (Thompson et al., 1999, 2007). Where exposed the intrusion has a V-shaped basal contact (inset B) and an erosive(scalloped) upper surface (C, D); for more detail see Scott et al. (2009). E) Irregular sandstone intrusions that display abrupt changes in thickness. F) An erosive convex-up upper contact. E and F are from the Marathon Basin, Texas (redrawnand modified from Diggs, 2007) and the lithologies represented are mudstone (dark grey) and sandstone (light grey).

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image of Fig.�10

Fig. 11. The external geometry and internal sedimentary structure of extrusive sandstones; extrusive and depositional sandstones are represented as light grey and dark grey,respectively. A) A plan view of a sandstone volcano with flow runnels radiating from a well-developed central vent. Sandstone volcanoes have a range of heights and widths. Widthranges between 0.3 and 1.8 m and height ranges 0.1 to 0.45 m (Ross Formation, County Clare, Ireland, redrawn and modified from Jonk et al., 2007). B) A plan view of a sandstonevolcano cluster. Sand volcanoes are up to 0.75 m in height and 3 m wide with a well-defined crater (black arrow). Flow runnels (white arrow) are present on the volcano flanks.Flanks are inclined at approximately 15° (Ross Formation, County Clare, Ireland, redrawn andmodified from Pringle et al., 2007). C) A tar-stained sandstone extrusionwith an overallfining-upward from coarse- to medium-grained sandstone with common mudstone clasts in the lowest 2 m of section to medium-grained sandstone with a sudden rapid reductionin grain size approximately 3.5 m from the base to siltstone. The siltstone is not tar saturated and contains abundant clasts of biosiliceous clay-sized material. The base of thesandstone is highly irregular with low-angle dikes running close to and disrupting the paleo-seafloor. High-angle dikes also disrupt the paleo-seafloor, typically forming sandstone-filled concave-upward ventsr. D) A bedding-concordant contact between a tar-stained sandstone extrusion and a paleo-seafloor in the Santa Cruz Mudstone. Note the absence ofdikes reaching the paleo-seafloor hence no disruption. Low-angle bedding is the prevalent internal structurewith localized soft-sedimentary deformation and bioturbation. The low-angle bedding defines low-relief conical mounds that are larger-scale analogues to the sand volcanoes in A and B. C and D are both from the Santa Cruz Mudstone at Red White andBlue Beach near Santa Cruz, California (see Hurst et al., 2006).


appear cross bedded (Boehm and Moore, 2002) and may reflectprolongedperiods of high-velocityflow fromamajor vent (Hurst, 2004)or interfingering and stacking of several volcanic cones. Preservation ofcentral vents in centers of sandstone volcanoes is unusual when theconical extrudite geometries overlap.

Numerous clasts of finer-grained host lithologies are locallycommon and are spatially associated with venting points (Obermeier,1996); clasts are interpreted to be derived from the underlyingsubstrate (Fig. 11C; Hurst et al., 2006). Normal gradation from basalclast-bearing sandstone to upper clast-free sandstone is common,

image of Fig.�11


which is interpreted to represent a transition from the violentturbulent eruption that accompanied the initial venting to ebbing ofthe flow as sand extrusion progressed (Obermeier, 1996). In othersandstone extrusions normal grading is crude (Hurst et al., 2006). Flowrunnels occur on the flanks of sandstone volcanoes and radiate awayfrom the crater (Jonk et al., 2007; Pringle et al., 2007; Fig. 11A, B).Bioturbation of extrusive sandstone is common (Fig. 11C,D), partic-ularly near the upper margin where contact with the overlyingsediment is often transitional (Hurst et al., 2006).

2.6. Parent units

Parent units are depositional sandstones that sourced sand andwere the reservoirs and conduits for the fluids that formed sandinjectites. Published information that describes the affect of post-depositional changes and sand mobilization on the geometry, internalstructures and microtexture of depositional sandstones is sparse(Appendix F). Most studies are confined to the subsurface, possiblybecause the external geometry of some large-scale remobilization ofdepositional sandstones is more readily identified using seismic data(Hurst et al., 2003a; Huuse et al., 2004; Szarawarska et al., 2010).

2.6.1. GeometryOutcrop examples of parent sandstone units that are modified by

sand fluidization and injection are shown in Fig. 12 butmost studies of

Fig. 12. Geometries in remobilized and injected depositional sandstones – parent units. SandA) Structureless sandstone with subhorizontal mudstone rafts of different shape and size; mufrom Surlyk et al., 2007, Figs. 13 and 14). B) Remobilized sandstone with an irregular upper mof the sandstone. Elongatemudstone clasts are foundwithin the sill (Jameson Land, Greenlandikes of medium-grained sand that crosscut the cross-bedded parent units. Internal sedimerosional and cut up to 0.1 m into the overlying strata (Charleston, South Carolina, redrawn aline) at the top of a depositional sandstone that truncates bedding in the overlying mudstonetar in the early 20th century (Santa Cruz Mudstone, Santa Cruz Co, California, modified from

the geometry of parent unit sandstones use subsurface data (Hurst etal., 2003a; Duranti and Hurst, 2004). Referring to outcrops Surlyk et al.(2007), recognize two separate groups of remobilized sandstonebodies: (1) laterally extensive and (2) mounded bodies (Hurst et al.,2003a). Laterally-extensive sandstones (Fig. 12A) commonly havesteps on both upper and lower margins. Step-up and step-down(often N10 m) are common and may cross 10 m or more hostmudstone before wedging out. Steps have a hierarchical arrangementwith large steps comprising smaller steps of 0.1 to N1 m in height.Mushroom-shaped geometries (Fig. 12B) are observed on the topsurfaces of laterally-extensive sandstone bodies that form highlyirregular geometries (Surlyk et al., 2007, Figs. 9 and 10). Thin sills(b0.2 m) commonly extend laterally from these irregular sandstones(Fig. 12A). Similar irregular upper surfaces on the tops of remobilizeddepositional sandstones are reported by Obermeier (1996; Fig. 12C).Convex-up erosive surfaces similar to those recorded on the uppersurfaces of sills (Figs. 5 and 6) and irregular sandstone intrusions(Fig. 10) are common andmay produce a dome ormounded geometry(Fig. 12).

Interpretation of seismic data may reveal the external geometry ofdepositional sandstones that were modified by sand injection (Hillierand Cosgrove, 2002) but external geometry alone is not definitiveevidence of a modified depositional origin and their interpretation isnot calibrated by core data. The paucity of relevant detailed outcropdata greatly limits the interpretation of subsurface examples. What is

stones and mudstones are represented as light grey and dark grey shades, respectively.dstone rafts can exceed 7 m in length (Jameson Land, Greenland, redrawn andmodifiedargin. The top margin displays concave-upward geometries. A sill extends from the topd, redrawn andmodified from Surlyk et al., 2007). C) Schematic vertical section showingentary structures are destroyed adjacent to the dikes. The upper surfaces of sills arendmodified from Obermeier, 1996). D) A convex-upward erosive surface (dashed black. This feature probably formed a mounded feature 20–30 m high that was excavated forHurst et al., 2007).

image of Fig.�12


clear from the published studies is that it is difficult to conceptualizehow fracturation of fine-grained host strata could have formed theirregular, often low-angle external geometries that are preserved.

2.6.2. Internal sedimentary structures and microtexturesMobilization of sand and fluids within and from parent units

induces changes to pre-existing depositional structures (Appendix F).The extent of sand mobilization and injection depends upon thephysical character (i.e. grain-size, volume of clay-size particles and

Fig. 13. Outcrop characteristics of a regionally-developed sand injectite, the Panoche GianA) Part of a regionally-developed sandstone intrusion with depositional strata dipping at ~30units outcrop in the SW (left) in which depositional sedimentary structures are preservfluidization and feature erosive (scalloped) surfaces, deformed depositional structures or tthem. The sill zone (Vigorito and Hurst, 2010) overlies the depositional sandstones, which inangle dikes bifurcate and narrow upward; their abundance and apertures are markedly lesseafloor and sandstone extrusions are absent in this view. B) Detailed view of the framed areand with the sills form a completely hydraulically-connected system of sandstones. BeddinGiant Injectite Complex (PGIC), California (Vigorito et al., 2008).

grain angularity) of sandstones prior to remobilization (Lowe, 1975)and the rate and duration of pore-fluid escape. A commonly-displayedfeature is the grading of pristine depositional structures into upward-deformed laminae, giant pillars, and structureless sandstones (Fig. 12;Hurst et al., 2003a; Duranti and Hurst, 2004). Banding (0.1 m scale)sub-parallel to the erosive upper margins is common (Fig. 12D). Thistransition records an accelerating fluidized flow as sand and fluidsmove upward and increasingly interact, reorganize and obliterate thedepositional fabric. Consequently, structureless sandstones often

t Injectite Complex (PGIC), California (Vigorito et al., 2008; Vigorito and Hurst, 2010).° to the NE (in the field of view the section youngs from left to right). Sandstone parented. Above this level depositional sandstones become increasingly modified by sandheir complete obliteration, and have numerous sandstone intrusions emanating fromturn is overlain by large low-angle dikes and lower aperture high-angle dikes. The high-s in and above the Marca Shale (pale-grey mudstones in right hand view). The paleo-a shown in A that focuses on the sill zone. Short, randomly-oriented dikes are commong in the mudstones is dipping to the NE at ~30°. From Dosados Canyon in the Panoche

image of Fig.�13


occur uppermost in a sequence of increasing intensity of fluidizationupward during a period of pervasive pore-fluid escape (Duranti andHurst, 2004).

2.7. Architecture of sand injectites

Sand injectites have a tri-partite organization: parent sandstoneunits that source sand and fluids for sand injection; an intrusivecomplex, that consists largely of dikes and sills; an extrudite complexthat is largely sandstone extrusions sourced from the underlyingintrusive complex (Vigorito et al., 2008). Since the outcrops of sandinjectites referred to in this paper are mostly small and/or relativelydiscontinuous, there is a significant limitation when using the outcropdata as analogies for those imaged in the subsurface. Only the PGIC(Panoche Giant Injectite Complex, Fig. 13) described partly by Smyersand Peterson (1971) and in more detailed summary by Vigorito et al.(2008), is known to have regional extent and contains sandstoneswith volumes and shapes that can be readily compared withsubsurface data. Outcrops with sand injectites that occur at a scalerelevant to that which can be imaged on seismic reflection data occurin the Hareelv Fm. (Surlyk and Noe-Nygaard, 2001; Surlyk et al., 2007)and to a lesser extent in the Miocene near Santa Cruz (Thompsonet al., 1999, 2007; Scott et al., 2009) but their limited regional extentand less constrained geological context limits their value as analogues.

Within the tri-partite organization in the PGIC distinctivesandstone architectural features occur (Fig. 13). Parent units have ageneral trend of increasingly deformed sandstones upward, caused bymore intense sand fluidisation upward, with high-angle, high-aperture dikes (Vigorito et al., 2008; Vigorito and Hurst, 2010). Theintrusive complex has a lower sill-zone the base of which is termedthe lithostatic equilibrium surface (LES, Fig. 14); this was an isobaric

Fig. 14. Schematic correlation between summary stratigraphic sections from the Panoche Giathe Panoche and Tumey hills. Note the variations in the sedimentology of the host strata as wand Hurst, 2010). The sill zone pervades the entire section and cuts lithostratigraphic and chAn unconformity (Eocene) erodes progressively deeper from north to south removing the

surface at the time of regional hydrofracture that caused sandinjection. The sill zone is remarkably uniform throughout ~400 km2

despite lithological variations in the host strata (Fig. 14; Vigorito andHurst, 2010). Sandstone intrusions in the sill zone lack consistentorientation and in detail they appear chaotic (Fig. 13B) however, sillstend to change geometry upward from staggered, to stepping andthen multi-layered geometry (Fig. 13). Above the sill zone dikesdominate with high-angle, bifurcating dikes becoming more abun-dant upward (Figs. 2A and 13A). The generic significance of the PGICsandstone architecture in other sand injectites is untested apart fromsome preliminary confirmation by Vigorito and Hurst (2010).

Studies of subsurface data, largelymotivated by the exploration forhydrocarbons, focus on sandstones that are large enough to containcommercial reserves of hydrocarbons and can be imaged usingseismic reflection data (summary in Huuse et al., 2007). As seismicreflection data rarely allow lithological units thinner than 5 m to beimaged and tend not to resolve the presence of steep features (e.g.sandstone dikes) there are immediate and obvious limits to howmuch of a subsurface sandstone injectite complex can be identifiedandmapped (Briedis et al., 2007). A further limitationwith subsurfacedata is that boreholes tend to sample the shallowest parts of sandinjectites because that is where migrating hydrocarbons will first betrapped. Hence the deeper parts of sand injectites, for example parentunits, may be significantly under-sampled by hydrocarbon-industryboreholes.

2.8. Host strata: sealing capacity

Sand injectites are commonly, although not exclusively (e.g. Netoff,2002; Huuse et al., 2005; Glennie and Hurst, 2007), found in host stratathat are fine grained and in general terms constitute aquitards or seal

nt Injectite Complex (PGIC) that runs approximately NNE (left) to SSW (right) throughell as variations in the architecture of the sandstone intrusions (modified from Vigoritoronostratigraphic boundaries (LES – lithostatic equilibrium surface, TSZ – top sill zone).top of the sand injectite.

image of Fig.�14


lithologies. To promote sand injection elevated pore-pressure isrequired (Jolly and Lonergan, 2002; Hurst et al., 2003a) and tomaintainelevated pore-pressure escape of the pressure requires a seal (aquitard)to reduce fluid flow toward the earth's surface. Although in widespreaduse in the oil industry, seal is probably a slightly misleading term as allnatural seals have an interlinked pore-system filled by formation fluids(Bjørkum et al., 1998). Hence seals are transient and inherently leaky.Consequently definitions of seal are ambiguous, some authors definingseal as a body of rock that is capable of preventing all pore-fluid flowthrough substantial periods of geologic time (Hunt, 1990) whereasothers recognize that pore entry and exit pressures control seal qualityat a granular scale, which in turn determine their permeability(Swarbrick and Osborne, 1998). It is unlikely that seals exist that havezero effective permeability to fluids over extensive periods of geologictime (Deming, 1994) and in the context of hydrocarbon systemsCartwright et al. (2007) use aworking definition of a “sealing sequence”as a low-permeability lithofacies that halts or retards the flow ofpetroleum toward the basin surface.

Following detailed outcrop-mapping a zone of regional hydraulicfailure was identified that corresponded to the occurrence ofhydrofractured mudstones and a concentration of sandstone sills(Vigorito and Hurst, 2010). This sill-rich zone (Fig. 14), the base ofwhich corresponds to an isobaric surface termed the lithostaticequilibrium surface has regional occurrence and above which porepressure is interpreted to have been supra-lithostatic at the time ofsand injection (Vigorito and Hurst, 2010). Location of the base of thesill zone is directly related to the thickness of the overburden at thetime of seal failure and sand injection. Although there is sporadicevidence (sills) for supra-lithostatic pore pressure below and, lessoften, above the sill zone in the PGIC (Vetel and Cartwright, 2009)these occurrences record localized not regional supra-lithostatic porepressure during sand injection.

Development of fluid overpressure is a balance between thepermeability of host strata and the rate at which pore fluid isintroduced into strata from the underburden. Hence if a charge of porefluid is rapid the permeability of the seal lithology is less influential onthe development of elevated pore pressure and cases exist wherelaminated fine-grained sand has been sufficiently low-permeability toact as a transient seal (Glennie and Hurst, 2007). In thick mudstonesspecific characteristics of the host strata (clay mineralogy, clay andorganic matter content, grain packing, pore-fluid properties) deter-mine the seal quality and in combination with burial history (rate ofburial, burial depths, temperature) where and when regionalhydrofracture will occur. Mud tends to undergo mechanical compac-tion during shallow (b1 km) burial, however, mechanical compactionis less important in deeply buried sediments, and temperature-dependent chemical processes exert more of a control on host-stratapermeability. This may explain the occurrence of fluid overpressureand the abnormal preservation of porosity in deeply buriedmudstones (N2 km) in the North Sea (Broichhausen et al., 2005).Modeling the chemical compaction of mudstones can generate pore-volume losses similar in magnitude to mechanical compaction(Matthews et al., 2000).

3. Sand injection: triggers, fracturation, and grain fluidisationand sedimentation

The general series of events that leads to the development of sandinjectites is well known, elevation of pore pressure, hydrofracture ofsealing strata, fluidization and injection of sand, and the origin of thearchitecture of regionally-developed sandstone intrusions can berelated to both the process of sand emplacement and the evolvingpore-pressure regime (Vigorito and Hurst, 2010). Precise details ofwhat triggers sand fluidization and the character of the flow processesresponsible for the development of discordant margins of sandstoneintrusions are in general poorly constrained. Equally, there are few

detailed analyses of the sedimentation processes responsible for theformation of the internal structures and microtextures.

3.1. Trigger mechanisms

A number of triggering mechanisms are associated with theformation of sand injectites including seismicity (Obermeier, 1996,1998; Rosseti, 1999; Boehm andMoore, 2002; Obermeier et al., 2005),overpressuring caused by rapid loading (Truswell, 1972; Allen, 1985;Strachan, 2002; Hildebrandt and Egenhoff, 2007), thermal pressuri-zation (Ujiie et al., 2007), and fluid migration (Jenkins, 1930; Davieset al., 2006). More unusual trigger mechanisms have been inferredfrom the geological record, such as impact cratering (Sturkell andOrmö, 1997; Alvarez et al., 1998) and igneous intrusion (Curtis andRiley, 2003). In a recent review seismicity and rapid loading wereconcluded to be the most typical triggering mechanisms (Jonk, 2010).

Seismicity is widely invoked as a trigger mechanism for sandinjection and recent extrusion of sand at the earth's surface related toearthquake activity is common (Obermeier, 1996; Obermeier et al.,2005). The spatial distribution of sand extrusions, soft-sedimentdeformation structures with faults and fractures are used to constrainthe epicenters of seismic activity (Obermeier, 1998; Rosseti, 1999;Rodriguez-Pascua et al., 2000; Obermeier et al., 2005) and highmagnitude (N6) earthquakes induce soft-sediment deformation 15–20 km from seismic epicenters (Rosseti, 1999). The frequency withwhich sandstone injectites are found in seismically-active areassuggests that it is a significant triggering mechanism (Saucier, 1989;Obermeier, 1996; Lawrence et al., 1999; Galli, 2000; Moretti, 2000;Boehm and Moore, 2002). However, if one considers the energyrequired to fluidize and inject the 10's of km3 of sand in regionally-developed sand injectites (Huuse et al., 2005; Vetel and Cartwright,2009; Vigorito and Hurst, 2010) earthquakes are unlikely to be thesole cause of sand injection (Duranti, 2007; Szarawarska, 2009).

Load-induced overpressuring is commonly caused by gravitationalinstability along submarine slopes that generates large-scale, down-ward mass-transport as sediment slides and slumps (Truswell, 1972;Strachan, 2002; Jonk et al., 2007), and storm waves that cause suddenloading of the sea floor (Hildebrandt and Egenhoff, 2007). Such eventstend to produce localized sand injection that is sometimes accompa-nied by sand extrusion, typically as sand volcanoes. Where sandextrusion occurs the timing of sand injection can be constrained(Hurst et al., 2006; Hildebrandt and Egenhoff, 2007). Mass-transportand storm-wave loading are associated with near sea-floor sandfluidization and injection (Fairbridge, 1946; Strachan, 2002; Hildeb-randt and Egenhoff, 2007) but sand injection related to load-inducedtectonism may occur deeper, probably at seismogenic depths of up to4–6 km (Ujiie et al., 2007).

Rapid migration of fluids into depositional sand bodies that causespore pressure to rise suddenly and thereby initiate sand fluidizationand injection has until recently been largely overlooked as a triggeringmechanism. Sourcing the rapid migration of pore water during burialmay have a variety of origins during shallow (b1.5 km) burial, forexample, formation of polygonal faults in mudstones (Cartwright andDewhurst, 1998; Cartwright et al., 2003; Wattrus et al., 2003),mineralogical phase changes, for example opal A to CT (Davies et al.,2006), and rapid migration of hydrocarbon gas (Brooke et al., 1995).Another major source of fluid to promote excess pore pressure couldbe the decomposition of gas hydrates during periods of eustatic sea-level change or ocean warming. All are mechanisms by whichsubstantial volumes of pore fluid are liberated from fine-grainedstrata and thereby may elevate pore pressure in sand/sandstonebodies and make them susceptible to sand fluidization and injection,assuming that the fluids released can migrate into appropriatestartigraphic levels. Large-scale sand injectites are partly attributedto the migration of fluids from deeper sedimentary sources and fluidmigration routes have been imaged on 3D seismic reflection data

Table 1Estimates of sand and fluid volumes during sand injection and fluid volumes liberatedby liquefaction and consolidation of parent sand units. VTI=gross volume of injectedsand at the time of injection (using Ø=60–80%). VT2=present day volume of injectedsand. Vg=grain (mineral) volume of injected sand body. Vfinj=fluid volume at the timeof injection. Vfcon=fluid volume at the present day. Volumes in m3. Data for the smalland large injectites are values taken from Injected Sands database (Universities ofAberdeen and Cardiff). Data for the consolidated parent unit is a maximum figure basedon the large composite consolidation complex in Fig. 2, Hurst and Cronin (2001). Tablecompiled and modified after Hurst et al. (2003a).

VTI VT2 Vg Vfinj Vfcon

Small injectite 6.5–7.5×104 5×104 3.5×104 3.0–4.0×104 1.5×104

Large injectite 5.2–6.0×107 4×107 2.8×107 2.4–3.2×107 1.2×107

Consolidated unit – 4×102 3.5×102 1.5×102 1.5×102


(Løseth et al., 2003; Hurst et al., 2003b; Huuse et al., 2005) and cancrosscut more than 200 m of fine-grained strata. Simple volumetriccalculations show that it is unlikely that the fluid volumes required tofluidize large sandstone intrusions of the type imaged on reflectionseismic data (MacLeod et al., 1999; Duranti et al., 2002a; De Boer et al.,2007) can be derived by liquefaction of the depositional sand bodiesfrom which they emanate (Table 1).

3.2. Fluid types

Occurrence of sand injectites in a stratigraphic interval implicitlysignifies a period of elevated pore-pressure, hydrofracture of a seallithology and fluid-flow vigorus enough to entrain sand and drive itinto a fracture system. Formation water is by far the most common

Fig. 15. Hydrofracture mechanisms. A) Mohr-failure envelope demonstrating the stress conB) The physical response of a material suffering extensional (red line) and shear failure (greefrom parent sand units. Hence the parent sand unit is unable to become overpressured. D) Hwill set up a pore-fluid pressure gradient from high to low pressure in the parent unit and ffluidize sand, into the fracture (modified after Cosgrove, 2001); σ1, σ2 and σ3 are the principlassumed to be vertical.

fluid in the shallow crust and is likely to be the main agent of sandfluidization however, hydrocarbons are considered as an alternativeor additional fluid (Jolly and Lonergan, 2002) and are detected in fluidinclusions in sandstone intrusions (Jonk et al., 2005a). Buoyancycreated by hydrocarbons seems unlikely to have a significant effect onpore pressure (Jonk, 2010) although the role of hydrocarbon gas as asupport to sand fluidization remains (Brooke et al., 1995; Hubbard etal., 2007), as does the generation of carbon dioxide and methane fromthe biodegradation of light hydrocarbon fluids.

3.3. Hydrofracturing

Independent of the triggermechanism seal failure is assumed to becaused by hydraulic fracture that that occurs when pore pressureexceeds (at least) the fracture gradient (Jolly and Lonergan, 2002)within the seal lithology or close to the contact between the seal andparent-unit sandstones. Dikes form both by tensile and shearhydrofracturing of host strata when the differential stress intersectsthe Griffith extensional failure envelope (red Mohr circle; Fig. 15A)and the Navier Couloumb failure envelope (green Mohr circle;Fig. 15A). Tensile (1) and shear fractures (2) are expressed by:

τ2 + 4 T σn−4 T = 0 ð1Þ

τ = C + μ σn ð2Þ

where τ is the shear stress, σn the normal stress, T is the tensile strengthof the rock, C the cohesive strength, and μ the coefficient of internalfriction. The relationship between the orientation of the principle

ditions for shear (green circle) and extension (red circle) of hydraulic-fracture failure.n line). C) Pre-existing fractures and low-capillary entry pressures allow fluids to escapeigh-pore-fluid pressures within the parent sand unit can initiate hydraulic fracture. Thisracture, respectively. Pore fluid will flow, and if the flow velocity is high enough it wille stress axes, maximum, intermediate and minimum, respectively – σ1 is by convention

image of Fig.�15


stresses and extensional and shear fractures show that failure occursunder differential stress (i.e. σ′1−σ′3, tensile failure; σ1−σ3, shearfailure). This situation can occur when the tectonic stress in a basin istensile, but also commonly occurs when the internal fluid pressure (p)in a sedimentary succession acts to oppose the applied stresses (σ1−p)(σ2−p) (σ3−p) (Cosgrove, 2001). In the context of sand injection thissituation occurswhen high pore-fluid pressure in a parent unit opposesthe principle stresses, thus causing the differential stress to intersect thehost strata's failure envelope (Cosgrove, 2001) and form an extensionalfracture (red box; Fig. 15B). High pore-fluid pressures reduce theeffective stress at grain-to-grain contacts and are expressed byTerzaghi's equation:

σp = po−pf ; ð3Þ

where σp is the effective stress carried by grains, po is the overburdenstress, and pf is the pore-fluid pressure. In general the mechanismsresponsible for pore-fluid overpressure are decreasing pore volume(disequilibrium decompaction, Swarbrick and Osborne, 1998) andincreasing fluid volume (Davies et al., 2006), or processes related todensity differences or fluid movements (for a review related to theseprocesses see Osborne and Swarbrick, 1997). Diagnostic features ofsandstone-filled hydraulic fractures include “chicken wire” (Powley,1990) and jigsaw textures (Duranti and Hurst, 2004).

Leakage of fluids from sandstones occurs when the capillaryentry pressure of the surrounding strata is lower than the fluidpressure (Fisher et al., 2001). In addition, sandstone may transmitfluids into lower-permeability host strata if, for example, there arefractures with sufficient permeability; this will dissipate thepotential build up of overpressure in sandstone (Fig. 15C). If theadjacent strata are of sufficiently low-permeability (i.e. seals) andthe pore-fluid pressure exceeds the fracture gradient of the seal, thehost strata will hydrofracture and will tend to form steep fracturesthat are align approximately parallel to the maximum and minimumhorizontal stress vectors (Fig. 15D). If during the propagation of ahydrofracture the differential pressure between the parent sand unitand the fracture is large, pore fluids will flow in accord with thepressure gradient. If the upward flow is greater than the minimum-fluidization velocity the parent sand unit will fluidize and inject sandinto the fracture (Fig. 15D). If pore pressure exceeds the lithostatic(overburden) gradient hydrofracture of a seal produces randomly-oriented fractures that include bedding parallel surfaces along whichsandstone sills may form (Vigorito and Hurst, 2010).

3.4. Transport and sedimentation: liquefaction and fluidization

Allen (1984) defines liquefaction as the breakdown of the fabric ofa material to such a degree that grains are no longer mutuallysupported but temporarily separated and dispersed in a pore fluid. Itinvolves the rapid rise of pore-fluid pressure and it can help promoteinjection of sand at shallow depths (b10 m, Obermeier, 1996).Fluidization is the suspension of grains by drag forces imparted byupward fluid-flow (DiFelice, 1995). It is themechanism bywhich sandis transported from parent sandstone units into hydrofractures(Duranti and Hurst, 2004).

3.4.1. LiquefactionLiquefaction occurs when grains are no longer supported by

intergranular friction but are momentarily suspended by the ambientpore fluid (Leeder, 1982). It tends to drive pore-fluid upward andinitiates localized sand fluidization (Obermeier, 1996; Hurst et al.,2003a). If flow occurs it is as a viscous fluid when deformed by shearstress caused by the reduced sediment strength (Allen, 1985; Owen,1996). During burial, incremental overburden-pressure increasessediment shear-strength that hinders liquefaction (Obermeier, 1996).

Mechanisms responsible for the upward movement of porefluids are: 1) high pore-fluid pressures in the liquefied sedimentthat drive fluids to the sediment surface and; 2) the resettlement ofgrains into more tightly-packed configurations that displace pore-fluids upward (Leeder, 1982). Sand grains fluidize when the upwardmovement of pore fluid is greater than their minimum-fluidizationvelocity. When liquefied sand is overlain by fine-grained low-permeability strata the rate of upward fluid-escape from the sand isslowed and excess fluid begins to accumulate below the low-permeability layer. If pore-fluid pressure exceeds the fracturegradient in the low-permeability layer hydrofracture of the layeroccurs, thus causing fluids and sand to vent at the surface(Obermeier, 1996). Pre-existing fractures may be importantinternal weaknesses that allow fluid and sand to vent more easilyat the surface. Fluidized sand forms fountain flows as high as 3–5 mabove ground level (Obermeier, 1996), which corresponds to exit-flow velocities of 1–10 ms−1 (Duranti and Hurst, 2004).

Liquefaction of sand is commonly caused by seismically-inducedshearing (Takahama et al., 2000) and associated with sand injectionand extrusion (Obermeier, 1996). Cyclic-shearing causes disaggre-gation of loosely-packed grains thereby transferring the support ofthe overburden pressure to the pore fluid. This causes the pore-fluidpressure to rise suddenly to reach the static confining pressure,which can initiate hydrofracture in the overlying sealing strata andsand injection may ensue. Pressure changes caused by loading fromlarge storm-waves (Hildebrandt and Egenhoff, 2007) or large, denseturbidity currents (Leeder, 1982) are inferred as possible mecha-nisms to induce liquefaction in underlying strata. Both thesemechanisms need not require an external source of fluid andrepresent in situ liquefaction within a closed reference volume(Owen, 1996).

3.4.2. FluidizationFluidization is regarded as the main process responsible for the

injection of sand into hydraulic fractures (Duranti and Hurst, 2004;Gallo and Woods, 2004). It forms by the suspension of grains by dragforces imparted by upward flow. It can occur either as particulate(homogeneous) or aggregative (inhomogeneous) fluidization. Com-plete fluidization occurs when all grains of different size and/ordensity are fluidized.

Sand injection is considered to form by the complete fluidizationand injection of sand into a fracture. In terms of grain-size, fine- tomedium-sand-size grains are preferentially fluidized because theyhave the lowest minimum-fluidization velocity (Lowe, 1975). Thismay help explain why the majority of injected sandstones arecomprised of fine- to medium-grained sandstone. Fluidizationrequires that there is an upward flow of fluid across granular material(DiFelice, 1995; Epstein, 2003; DiFelice, 2010). An overpressuredsandstone unit will fluidize when the pressure gradient is sufficient todrive upward-flow above the minimum-fluidization velocity (Vmf)(Gallo and Woods, 2004). Typically, the Vmf is estimated (after Lowe,1975) as:

Vmf = 0:00081 ρs−ρfð Þð Þd2g = μ ð4Þ

where ρs=density of solid, ρf=density of solid, d=mean diameter ofparticles (mm), gravity and μ=viscosity. The Vmf for fine- to medium-grained sand is estimated to be as low as 0.001 ms−1 (Lowe, 1975)and 0.01 ms−1 (Duranti and Hurst, 2004). These are minimum flowvelocities when fluidization begins and may represent the flowvelocity for small-scale dewatering structures such as pipes (Lowe,1975). Large-scale sand injection (e.g. kilometer-scale conical intru-sions, Huuse et al., 2005) is likely to be produced at much higherfluidization velocities. Commonly, fluidization velocity during sandinjection is estimated using the square-root law, which is suitable forgrains over a fewmillimeters in diameter, from the minimum-settling


velocity (Vt) of large clasts of mudstone found within injectedsandstone (cf. mudstone-clast breccias, Duranti and Hurst, 2004), as:

Vt = k ρs−ρf� �

= ρf� �

ÞgdÞ1=2; ð5Þ

where Vt=minimum-settling velocity (m s−1), ρs=density of themudstone clast (kg m−3), ρf=density of fluid (kg m−3), g=accel-eration due to gravity (9.81 ms−2), d=mudstone clast diameter (m),and k=is a constant, where: k=(4/3 CD, 0)1/2; the drag coefficient fora solitary sphere in approximate motion (CD, 0) is constant at 0.45over a wide range of Reynolds numbers (Re) (i.e. from 1×103 to3×105) (Allen, 1985). Vt is estimated to be many orders ofmagnitudes greater than Vmf (0.1–0.5 ms−1, Duranti and Hurst,2004; 0.7 ms−1, Hubbard et al., 2007; Gamberi, 2010).

3.4.2.1. Fluidized flow during sand injection. Sand injection occurswhen sand-charged fluids emplace sand into a fracture network inadjacent, typically overlying, low-permeability strata (Duranti andHurst, 2004). Some evidence of the flow processes during sandinjection is derived from examining the character of the discordantmargins between sandstone intrusions and their host strata (Diggs,2007; Hubbard et al., 2007). Precise processes that develop thediscordances, such as flutes, grooves, rills, and gutter casts, are poorlyconstrained but are inferred to be caused by erosion (Archer, 1984;Kawakami and Kawamura, 2002; Diggs, 2007). In the absence ofgeological experiments that study sand fluidization, experiments thatstudy fluidized granular flow in pipelines are used as analogues forfluidized flow in fractures.

Physical experiments demonstrate that the transport of fluidizedgrains causes the erosion of (fabricated) pipe margins (Mills et al.,2004) and the degradation of grains (Werther and Reppenhagen,2003). The erosive processes commonly responsible for erosionalmargins and grain degradation are abrasion (Deaquino et al., 1998)and corrasion (Werther and Reppenhagen, 2003); these form part of acontinuous spectrum of erosive processes collectively termed attri-tion. Erosion of the pipe margins and grain degradation are mostsubstantial at high fluidized-flow velocities (Werther and Reppenha-gen, 2003; Mills et al., 2004). Other erosive mechanisms that are notreplicated in experiments occur in nature and are inferred from therock record such as frictional drag, which may be responsible forforming some discordant marks (Diggs, 2007).

In the light of the experimental studies the erosion of fracturemargins during sand injection is entirely reasonable and is suggestiveof the erosive interaction between high-velocity fluidized sand andthe host strata. Erosional margins occur as scallops that are knownfrom outcrop studies to erode up to ~30 m into the host strata(Duranti et al., 2002b) more typically up to ~5 m Vigorito et al., 2008;Scott et al., 2009). This scale of erosion along the injection margins islikely to record sustained high-velocity fluidized-flow. Formation offlutes is by analogy with depositional environments direct evidence ofturbulent flow.

Degradation of grains, particularly of mudstone clasts, is acommon microtextural characteristic of sandstone intrusions. Severalintegrated outcrop and petrographic studies describe mudstone clastswith erosional margins and layers of embedded sand grains withinthe clasts (Kawakami and Kawamura, 2002; Scott et al., 2009). Scott etal. (2009) argue that these textures are likely to have formed bycorrasion, a process whereby high-velocity grains impact and embedinto (mudstone) clasts at non-zero angles (Allen, 1984, 1985;Werther and Reppenhagen, 2003).

3.4.3. Sedimentation during sand injectionSedimentary structures and their microtextural character provide a

record of theflowand sedimentation processes during sand injection. Incommonwithmany depositional environments sedimentary structures

may record only the final stages of deposition during sand injection. Insandstone intrusions depositional sedimentary structures are usuallyonly post dated by elutriation structures, typically pipes (Scott et al.,2009). During the injection of sand-charged fluids it is unknownwhether the grains settle onto fluid-sediment interfaces and move asbedload layers in traction, or whether they move as cohesive plug-likeflows that settle en masse, or as a combination of these processes. Thelack of consensus is due, in part, to the lack of outcrop studies that havefocused on the occurrence and character of sedimentary structures, butis also because of a lack of experimental data that capture thehydrodynamics of sedimentation during sand injection. Until the linksbetween the style of sedimentation and sedimentary structures are fullyunderstood there will be limits to robustness of inferences about flowprocesses inferred from the rock record.

Experimental data on the sedimentation of fluidized grains insettling columns help to infer processes that are likely to beresponsible for sedimentary structures observed in sandstone intru-sions; grading has been reproduced experimentally during thesedimentation of fluidized grains at different grain concentrations(Druitt, 1995). Druitt's study shows that above a critical concentration(45 to N60% depending on the grain size distribution) grain-sizesegregation during sedimentation is hindered and deposits with poorgrain-sorting form. It is thus reasonable to infer from outcrop studiesthat report grading in sand injectites, which includes spectacularexamples of normal grading in sandstone dikes (Hubbard et al., 2007),that the grains settled in a relatively low grain-concentration fluid (45to N60%, Druitt, 1995).

Although the experimental studies help to constrain the process ofsedimentation from fluidized flows they consider only the settling ofgrains in a stationary fluid (i.e. horizontal flow). The commonoccurrence of laminae and imbricate clasts in injected sandstones(Kawakami and Kawamura, 2002; Scott et al., 2009) suggests thatbedload layers moved during sustained periods of traction (i.e.evidence of horizontal-flow). Unfortunately there are no experimentsthat focus on fluidized flow in fractures; thus the precise flowconditions conducive to the transport of grains in bedload traction infractures during injection are unknown. Nevertheless it is possible toplace some constraint on these processes by considering the experi-mental studies on grain flow; in particular, when fluidized grains areforced to flow as sand-charged fluids along horizontal pipelines.

Laboratory experiments (Lee et al., 2004) demonstrate that thehorizontal transport-mechanism of grains depends upon the Froude(ratio of inertial forces to gravitational forces) and Reynolds (ratio ofinertial forces to viscous forces) numbers. Grains move in bedloadtraction below critical Reynolds and Froude numbers whereas abovethis threshold the grains move in suspension. Bedload layers movingin traction tend occur at low flow-velocities when the flow is belowthe critical “pick-up” velocity required to carry the grains insuspension (Lee et al., 2004). Therefore the occurrence of laminaesuggests that at the time of sedimentation during sand injection thefluidized flow-velocity was insufficient to keep the grains insuspension.

3.5. Fluidized-flow regime during sand injection

There is little geological consensus on the fluidized flow-regimeduring sand injection and it is attributed to both laminar (Dott, 1966;Peterson, 1968; Sturkell and Ormö, 1997) and turbulent flow(Obermeier, 1996; Duranti, 2007; Hubbard et al., 2007). The lack ofconsensus is demonstrated by contrasting interpretations of thesignificance of grain-size variation in sandstone dikes, whichwas usedto infer that normal grading in clastic dikes was indicative of theimportance of fluid turbulence during sand injection (Hubbard et al.,2007) but conversely, was used as evidence of a viscous-laminar flowregime because of the presence of graded layering (Peterson, 1968).


In order to resolve the ambiguity regarding flow regime duringsand injection some authors have estimated the Reynold's Number(Re),

Re = Vf a ρf� �

= μ ð6Þ

where Vf=fluidization velocity (ms−1), a=fracture aperture (m),ρf=density of fluid (kgm−3) and μ=fluid viscosity (Pas). Duranti(2007) modeled energy and flow velocities that were required forlarge-scale sand injection in the Paleogene of the North Sea andcalculated Re=1.4×104 to 1.4×105 values that are greatly in excessof the critical flow transition from laminar to turbulent flow infractures (~2300, Singhal and Gupta, 1999). The results are consistentwith more recent estimates of flow regime during sand injection(Re≈2.1×106, Scott et al., 2009).

3.6. Process of sand injection: overpressure development, liquefaction,hydrofracturing and fluidization

Fluid overpressure is a requisite precondition for sand injection tooccur as it initiates hydrofracture and drives subsequent fluid flow.Overpressure development is commonly assigned to disequilibriumcompaction during burial but other mechanisms, such as the injectionof fluids into a depositional sandstone and load-induced overpressur-ing, may demonstrate coupled relationships between overpressuredevelopment and it's dissipation during seal hydrofracture and sandfluidization and injection.

Sandstones become overpressured when the rate of compaction-induced fluid expulsion is reduced by low-permeability strata; once asandstone is effectively sealed (X1, Fig. 16) pore-fluid pressure willrise above the hydrostatic pressure-gradient (X1−X2, Fig. 16). In thisphysical state pore-fluid pressure resists mechanical compaction bythe overburden and thus maintains a metastable grain fabriccharacterized by high porosity (stage 2, Fig. 16). The pore-fluid

Fig. 16. Summary of the relationship between overpressure development with increasing burliquefaction. A) A possible pressure-depth path (dashed black line) for a sand body endisequilibrium compaction (X1–X2). At X2 shear-induced liquefaction rapidly raises the porinitiates hydrofracture, which creates a pressure gradient and in response, rapid pore-fluid echanges that accompany overpressuring in a sand body. Pore-fluid overpressure generatedporosity by transferring the load pressure to the pore fluid and thus resisting mechanical comgrain-to-grain contacts; momentarily the overburden is no longer supported by the fabricpressure (X3). Propagation of the hydrofracture establishes a pressure gradient with the oveinjects fluidized sand into the fractures. Resettling of grains intomore tightly-packed configu

pressure could in theory continue to increase by disequilibriumcompaction until it reaches the fracture pressure and initiateshydrofracture and sand injection. However, the slow rate of porepressure elevation associated with disequilibrium compaction requireexceptional low rates of leakage in seals, which is rarely the casewhenheterogeneities such as grain size variations and microfractures occurthat locally enhance permeability and thereby allow the dissipation ofoverpressure (Duranti and Hurst, 2004),. It is likely that the suddenfailure of thick sequences of low-permeability host strata associatedwith large sand injectites requires a trigger mechanism that canrapidly raise the fluid pressure to exceed the fracture pressuregradient and cause hydrofracxture.

A record of regionally-persistent supralithostatic pore-pressureduring sand injection is preserved throughout the Panoche GiantInjectite Complex (PGIC) independent of lateral variations in thedepositional host-strata (Vigorito and Hurst, 2010). This interval,termed the sill zone (Fig. 14), in which sills are common, dikes areshort and randomly-oriented, and mudstones contain hydrofractures(Fig. 13B), formed in response to supra-lithostatic hydrofracture.Above and below the sill zone dikes predominate and formed wherepore-pressure exceeded the fracture gradient. Formations of sillsrequire that the overburden is lifted as the sills form although thethickness of sills is augmented by upward erosion of mudstone andpossibly aided by the presence of bedding heterogeneities. Formationof isolated sills is a record of localized supra-lithostatic pressurewithin intervals where otherwise the fracture gradient is exceededand dikes form. A regionally-developed sill zone implies a regional-development of supra-lithostatic pressure.

A rapid increase in the pore-fluid pressure could be caused byshear-induced liquefaction related to seismicity although othermechanisms are possible such as sediment loading, thermal pressur-ization and injection of fluids. Liquefaction processes are believed tobe limited to shallow depths (b10 m). At greater depths, overburdenincreases sediment shear-strength thereby hindering liquefaction(Obermeier, 1996). If the sandstone unit is already overpressured (i.e.

ial depth due to disequilibrium compaction, and initiation of injection by shear-inducedveloped in low-permeability strata. Once enveloped (X1) the sand body undergoese-fluid pressure (X2–X3) and at X3 the pore pressure equals the fracture gradient; thisscape drives sand fluidization and injects sand into the fracture system. B) Grain fabricby disequilibrium compaction (X1–X2) preserves a metastable grain fabric with highpaction. Shear-induced liquefaction causes a breakdown of the grain fabric by removing, but by the pore fluid, which rapidly raises the fluid pressure (X2–X3) to the fracturerpressured sand body, which drives flow above the minimum-fluidization velocity andrations during the deceleration of fluid flow further helps to drive flow into the fractures.

image of Fig.�16


X2, Fig. 16) a metastable grain fabric is expected (stage 2, Fig. 16),which causes the grain contacts to break down under shear (stage 3,Fig. 16); this transfers grain support of the overburden to the porefluid, which rapidly raises the pore-fluid pressure (X2−X3, Fig. 16) andliquefies the sand. If the pore-fluid pressure reaches the fracturepressure (X3, Fig. 16 this initiates hydrofracture propagation, whichestablishes a pressure gradient between the sandstone unit and thefractures and drives fluidized sand into the fractures (stage 4, Fig. 16).

Of course this is a simple end-member hydrodynamic scenario thatcan lead to sand injection, and liquefaction alone is unable to accountfor large-scale remobilization and injection of a sand body because theliberated fluid volumes are insufficient to sustain fluidization (Hurstet al., 2003a). External sources of fluid are necessary. The volume offluid expelled by liquefaction and consolidation of depositional sandsand the volumes responsible for small and large-scale sand injectioncan be estimated by inferring rock and fluid characteristics at the timeof injection (Hurst et al., 2003a). Fluid volumes liberated duringliquefaction, estimated from the volumetric abundance of dish andconsolidation laminae (see Table 1) are two to five orders ofmagnitude less than that estimated for small (5×104) and large-scale injection (4×107). The scarcity of dish structures and consol-idation laminae in sandstones associated with sand remobilization(Dixon et al., 1995; Hurst et al., 2003a) and the lack of crosscuttingrelationships among the water-escape structures (e.g. Duranti andHurst, 2004), suggest that simultaneous liquefaction events togenerate enough fluids for large-scale injection are unlikely (Hurstet al., 2003a). Therefore it is concluded that liquefaction maycontribute fluids but volumetrically these fluids are insufficient tocause large-scale sand remobilization and injection. Rather, thegeneration of large-scale injection requires a combination of favorablefactors, such as overpressuring of depositional sandstone units causedby lateral compression at destructive plate margins (Ujiie et al., 2007),together with more rapid mechanisms of overpressure generation,such as the rapid introduction of fluid to raise the pore pressure abovethe fracture pressure and induce large-scale sand injection. Ulti-mately, the scale of sand remobilization and injection is determinedby the mechanisms and volumes of overpressure generation which inturn depends on the burial history of basins, including sedimentationrates, tectonism, diagenesis and other factors (Osborne and Swar-brick, 1997).

4. Post-sand-injection fluid flow and diagenesis

Immediate post-emplacement fluid flow in sandstone intrusions isrecorded by the common occurrence in dikes and sills of elutriationstructures, in particular pipes (Fig. 4D; Scott et al., 2009). Pipes formwhenmetastable granular fabrics collapse (Lowe and LoPiccolo, 1974)and in sandstone intrusions they cross cut pre-existing internalstructures and are likely to locally enhance vertical permeability(Hurst and Buller, 1984). Outcrop and petrographic data indicate thatsandstone intrusions act as high-permeability conduits for themigration of fluids, as is evidenced by the presence of residualhydrocarbons (Jenkins, 1930; Jonk et al., 2003; Mazzini et al., 2003a).Petrographic and fluid-inclusion data give further confirmation oftheir action as fluid conduits (Jonk et al., 2003, 2005a,b). Cathodolu-minescence images of carbonate cements in some sandstone injectiteshave a single cement phase, which supports that they were short-lived flow conduits (Mazzini et al., 2003b).

5. Petrophysical characteristics

Sand injectites have a broad significance to hydrocarbon explora-tion and production (Hurst and Cartwright, 2007) despite which thereare limited published porosity and permeability data, the twocharacteristics that impart primary control on the storage and flowcapacity of rocks, respectively. Some subsurface (borehole) data are

reported (Duranti et al., 2002a; Duranti and Hurst, 2004) but there areno reports of outcrop data.

Sandstone intrusions from Alba field (UK) have lower porositiesthan their parent sandstone units Duranti et al. (2002a) and otherstudies show that, on average, porosity is 2–3% lower in sandstoneintrusions relative to sandstones in parent units (Fig. 17). Duranti etal. (2002a) and Duranti and Hurst (2004) concluded that the lowerporosity is mainly caused by tighter grain packing (Hurst and Cronin,2001) and a slightly higher interstitial clay content. The enrichment ofinterstitial clay was attributed to the elutriation of clay-size particlesderived from the parent beds during the late stages of sand injection(Lonergan et al. (2007) attributed the lower porosities to theoccurrence of host mudstone clasts and deformation bands.

In general the porosity distribution of sandstone intrusions fromthe subsurface examples is broader than in parent units (Duranti et al.,2002a; Duranti and Hurst, 2004; Briedis et al., 2007; Lonergan et al.,2007). This broad distribution is thought to reflect the wide range ofgrain textures and grain packing caused by fluidization (Duranti et al.,2002a). No indication is given of the processes responsible forproducing this broad range. In particular the processes responsiblefor causing the variation in grain packing, which occur during sandfluidization, are not understood. Furthermore the spatial distributionof porosity within remobilized sandstones and intrusive elements isnot documented; this is largely due to the fact that petrophysicalstudies have focused on core data that provides little informationabout lateral trends in porosity and permeability, which are criticalwhen establishing their spatial distribution.

Sandstone intrusions are regarded as high-permeability conduits(Fig. 18) that can compromise the seal integrity of thick low-permeability mudstones (Løseth et al., 2003) and their ability toenhance the vertical permeability of a fine-grained sedimentarysequence can increase the rate of fluid migration within petroleumbasins (Hurst et al., 2003b). Sandstone intrusions are capable ofaltering the field-wide performance of reservoirs (Briedis et al., 2007;Fretwell et al., 2007; Guargena et al., 2007) and they can createsignificant cross flow between reservoir units that are separated byregionally-developed seals (Satur and Hurst, 2007).

To date, no outcrop studies of permeability are published and onlytwo core-based studies have examined the permeability character-istics of injected sandstones (Duranti et al., 2002a,b; Briedis et al.,2007). The mean permeability of sandstone intrusions shows only aslight reduction when compared to the parent units. Interestinglywhen compared with porosity data, the permeability distribution inparent units is broader than in the sandstone intrusions. Duranti et al.(2002a) attributed this to the process of sand remobilization, where inthe parent units bedding surfaces and laminae result in a heteroge-neous permeability structure, but the obliteration of these internalsedimentary structures during sand fluidization produces a moreisotropic and homogeneous microtexture. This is perhaps surprisingwhen considering the numerous internal sedimentary structures (i.e.laminae, graded layering, mudstone-clast breccias, and banding) havebeen reported from sandstone intrusions (Kawakami and Kawamura,2002; Diggs, 2007; Scott et al., 2009).

6. Discussion

6.1. Overview

Given the almost two-hundred year longevity of reporting thepresence of sand injectites (at least sandstone dikes; Murchison,1827) in the geological literature it is remarkable that theirdocumentation was sparse until this millennium (e.g. (Peterson,1968; Smyers and Peterson, 1971; Truswell, 1972; Lewis, 1973;Hiscott, 1979; Taylor, 1982; Martill and Hudson, 1989). Perhapsequally remarkable is the lack of realization that sandstone intrusionsrecord rather significant periods in the evolution of the shallow crust

Fig. 17. Frequency distribution of porosity (%) in parent sandstone units affected by sand fluidization and sandstone intrusions. In all cases the sandstone intrusions have lowermeanporosity (%) than their respective parent sandstone units; however the porosity of the sandstone intrusions remains high. Solid and dashed lines represent the parent sandstoneunits and sandstone intrusions, respectively. Porosity data from: the Nauchlan Member of the Alba Formation, Alba Field – blue (from Duranti et al., 2002a) and red (from Durantiand Hurst, 2004); the Lista, Sele, and Balder formations, Balder Field – green (after Briedis et al., 2007); wells in the Balder and Horda formations, Gryphon Field – pale orange andpurple lines (after Lonergan et al., 2007).


when regional excess pore-pressure developed along with large-scalehydrofracture and fluidization andmobilization of 10's to 100's km3 ofsand. Formation of sandstone intrusions and specifically dikes by thismechanism is so obviously different to the formation of Neptuniandikes (a sedimentary dike formed by infilling of sediment, usuallysand, in an undersea fissure of hollow; Bates and Jackson, 1980 –

subaerial examples also exist e.g. Simms, 1994; Di Stefano and

Fig. 18. Frequency distribution of permeability (mD) in parent sandstone units affectedby sand fluidization and sandstone intrusions. Sandstone intrusions have a lower meanpermeability (av. 2500 mD, Briedis et al., 2007; av. 2300 mD, Duranti et al., 2002a) thantheir respective parent sandstone units (av. 4000 mD, Briedis et al., 2007; av. 3300 mD,Duranti et al., 2002a). Solid and dashed lines represent the remobilized parentsandstone units and sandstone intrusions, respectively.

Mindszenty, 2000) – as to make any confusion between them likelyonly to be a function of the availability of too little diagnostic data.

When petroleum geologists working with borehole data (Dixon etal., 1995) recognized that substantial reservoir volumes weresandstone intrusions rather than depositional sandstones theircommercial significance was apparent. Appreciation of the volumeand distribution of sandstone intrusions, and their prospective andcommercial siginificance to the oil industry, did not become fullyapparent until the interpretation of reflection seismic data calibratedwith borehole data confirmed the presence of large, sandstone-rich,low-angle dikes (MacLeod et al., 1999; Duranti et al., 2002a). Thissubsequently led to the mapping of regionally-developed sandinjectites (Shoulders and Cartwright, 2004; Huuse et al., 2005), thesuccessful exploration for oil in part of a regional sand injectite(Rawlinson et al., 2005; De Boer et al., 2007), and the definition of anew trapping style, intrusive traps (Hurst et al., 2005). Subsurfacestudies demonstrated that very significant volumes of sand aremobilized, often on a basin-wide scale during sand injection (Hurstet al., 2003a; Løseth et al., 2003; Huuse et al., 2005), which until recentwork on thePanocheGiant Injectite Complex (Vigorito et al., 2008; VetelandCartwright, 2009;Vigorito andHurst, 2010)hadnot been recognizedin outcrop studies.

Despite the documentation of sand injectites in subsurface studiesthere may be some reluctance to accept that they are anything morethan “unusual features” that have some significance in oil and gasfields of the North Sea. The reluctance may be based on scientificreasoning but in the absence of any published evidence for this it ismore likely to be because sand injectites both as exploration targetsand in reservoir models is considered to be insignificant. Even giventhe limitations of the data referred to in this review and in Hurst andCartwright (2007a) the evidence points to sand injectites havingglobal occurrence throughout geologic time (Hurst and Cartwright,2007; Huuse et al., 2010). Sand injectites have a long pedigree in thestudy of deep-water clastic environments (Jolly and Lonergan, 2002;Hurst et al., 2007) but we observe surprisingly little transfer of ideaswithin or from petroleum systems where sand injectites are known tooccur (North Sea, offshore Angola, San Joaquin Valley) that may

image of Fig.�17

image of Fig.�18


increase recognition or even more testing of their commercial value.More fundamentally, as histories of elevated pore-pressure arefundamental in the formation of sand injectites (Vigorito and Hurst,2010) failing to include their occurrence and role in basin analysisignores sand injectites as a record of the evolution of pore pressureand basin-scale fluid migration.

It is likely that the lack of studies of high-quality outcrop of sandinjectites has limited their adoption into academic and appliedgeoscience groups. Detailed outcrop descriptions are published (forexample, papers in Hurst and Cartwright (2007) but few are made onextensive areas of high-quality exposure (Hurst et al., 2007). Conse-quently, when faced with examples of small sandstone intrusions mostgeologists have little field of reference or conceptual models that theycan use to create hypotheses about the likely geometry and significanceof the features that they are observing.

There has been a marked shift from observations of sand injectitesin subsurface data (since Dixon et al., 1995 to the present) to anincreased awareness of their more common occurrence at outcrop(Surlyk and Noe-Nygaard, 2001; Jolly and Lonergan, 2002; Netoff,2002; Parize and Friès, 2003; Parize et al., 2007a). More recently,study of the regionally-developed Panoche Giant Injectite Complex(PGIC) sand injectite at outcrop (Vigorito et al., 2008) allowedconsideration of the emplacement mechanics (Vetel and Cartwright,2009) and pore-pressure evolution (Vigorito and Hurst, 2010) of sandinjection at a scale similar to that known to exist in the subsurface(Huuse et al., 2004; Shoulders and Cartwright, 2004). Detail from thePGIC studies gives an opportunity to re evaluate and refine subsurfacestudies that are based largely on the interpretation of seismicreflection data, which typically cannot, for example, differentiatebetween erosive surfaces and a series of sub-decameter steps alongthe margins of sandstone intrusions (Fig. 6D) simply because oflimited resolution (Huuse et al., 2007). The excellent quality andextent of the PGIC outcrop has allowed an improved understanding ofthe distribution and significance of sills that occur pervasively abovelithostatic equilibrium in response to hydrofractiure of mudstones atsupra-lithostatic pore pressure (Vigorito and Hurst, 2010). Therelationship between in situ stress at the time of sand injection anddike orientation and spatial distribution is possible to study althoughin recent work it is unclear whether dike distribution in the PGIC israndom (Vetel and Cartwright, 2009) or more predictable (Vigoritoand Hurst, 2010) as was concluded earlier (Smyers and Peterson,1971). As the PGIC is the only regionally-developed sand injectiteoutcrop of sand injectite this places some limitations on its value as adirect analogue to other sand injectites however, if the relationshipsdescribed there are generic they have global applicability.

6.2. Identification of sand injectites

We suspect that many sandstone intrusions are misinterpreted asdepositional sandstones in outcrop and in subsurface studies.Spectacular subsurface examples of commercial significance are thedevelopment of the Balder, Alba and Gryphon oilfields, North Sea inwhich the recognition of the importance of sand injectites changedgeological models from predominantly depositional to increasinglysand injectite during 30, 15 and 10 year periods, respectively (Hurst etal., 2005). Similar outcrop examples where recognition of the role ofsand injection has increased through time are the interpretation of theHareelv Formation (compare Surlyk, 1987 with Surlyk et al., 2007)and the Pil'sk Suite (Macdonald and Flecker, 2007). Discordantmargins (Figs. 3, 6, 7A, 10 and 12) remain an important characteristicby which sandstone intrusions can be differentiated from depositionalsandstones. Although steep- and shallow-angle discordances withbedding are observed commonly from seismic reflection data(Duranti et al., 2002a; Huuse et al., 2005; De Boer et al., 2007), inboreholes (Purvis et al., 2002; Duranti and Hurst, 2004) and at outcrop(Parize and Friès, 2003; Diggs, 2007; Scott et al., 2009) alternative

interpretations may be invoked. At seismic scale individual saucer-shaped sills and conical intrusions of sandstone have been interpretedas cross sections of channels a hypothesis that can usually bedismissed when features are observed in plan-view as channelstend to have elongate, sinuous geometry and sand injectites do not(Huuse et al., 2005, 2007).

When viewed in core or in discontinuous outcrop some lithofaciestypical of sand injectites may appear similar to depositionalsandstones. Two example lithofacies that cause confusion aremudstone-clast conglomerates and structureless sandstones. Angularmudstone-clast conglomerates (also termed shale- or mudstone-clastbreccias because of their clast angularity) are common in sandinjectites (Figs. 6, 8 and10: Duranti et al., 2002a; Duranti and Hurst,2004;) but are common also in some depositional environments, forexample, as bank-collapse facies along fluvial or deep-water clasticchannel margins or as sandy debrites; small-scale characteristics ofthe mudstone-clast conglomerates often provide critical insight.Microfractures, often sandstone-filled, in mudstone clasts (Fig. 8)are formed by hydrofracturing and are diagnostic of sand injectites(Kawakami and Kawamura, 2002; Duranti and Hurst, 2004; Diggs,2007), are often associated with “jigsaw” clast-fabrics (Duranti andHurst, 2004) and may form in sand (Hurst and Glennie, 2008).Mudstone-clasts in depositional environments form predominantlyby gravitationally-controlled disaggregation which is unrelated tohydrofracturing.

Structureless sandstones are associated with rapid deposition ofhigh- to medium-density turbidites however in sand injectites theytypify sandstones in which intense and pervasive sand fluidizationoccurred (Duranti and Hurst, 2004). When viewed in core it isextremely challenging to differentiate between structureless sand-stones formed by rapid deposition and those formed by fluidizationand their larger-scale associations may provide the only clues to theirorigin.

Recognition of architectural elements in sand injectites is so farlimited to work on the Panoche Giant Injectite Complex (PGIC)(Vigorito et al., 2008) and brief comparison with four other sandinjectites (Vigorito and Hurst, 2010, Fig. 2). More data are requiredto substantiate if the tri-partite organization of parent sandstoneunits, an intrusive complex and an extrusive complex (extrudites) iscommon to all sand injectites. If an intrusive complex and a sill zone(Figs. 1 and 13) are identified inferences about spatial andgeometrical features of sandstones are possible as well as estima-tion of pore-pressure evolution during sand injection. Absence of asill zone implies that the pore pressure exceeded the fracturegradient in the overlying mudstones but did not exceed lithostaticpressure and that the sand injectite will be dominated by dikes.Comparison between the PGIC and Paleogene injectites in the NorthSea subsurface reveals that the PGIC intruded into a mudstoneoverburden that was twice as thick as the North Sea; the sandstonecontent of the PGIC is approximately half that of the North SeaPaleogene sand injectites.

6.3. Injectite margins and internal structures

Erosional margins are typical of sandstone intrusions (Figs. 6, 7and 10) and upward erosive margins characterize sandstone sills(Duranti et al., 2002a,b; Obermeier, 1996; Kawakami and Kawamura,2002; Macdonald and Flecker, 2007) and sandstone intrusions withmore irregular geometry (Hamberg et al., 2007; Scott et al., 2009).Upward erosive margins of sills help with the differentiation betweenthem and depositional sandstones in which downward erosivefeatures are pervasive. Identification of upward erosion may beparticularly useful in borehole studies where only narrow sections ofrock are sampled (typically b0.2 m) and in which bedding-concor-dance tends to dominate. Erosion in sand injectites is inferred to beindicative of high-velocity flow during injection (Vigorito et al., 2008;


Scott et al., 2009) and structures formed along margins such as flutesare indicative of turbulent, dilute flow, probably at high velocity(Hubbard et al., 2007). The material properties of fine-grained hoststrata during sand injection are however largely unknown and arelikely to affect the susceptibility to erosion.

Experiments indicate that the composition of fine-grained sedi-ment, such as clay fraction and water content influence erosion rates(Allen, 1984, 1985) and since these properties are strongly influencedby burial-induced compaction (Velde, 1996), then the depth of burialat the time of injection is likely influence the susceptibility to erosion.Most sand injectites appear to have formed at b1.5 kmburial (Jolly andLonergan, 2002) at which depth themudstone host stratawill often bepoorly consolidated (Velde, 1996). Despite this the predominant styleof mudstone fractures, including the geometry of detached fragmentsof host mudstone, is angular (Figs. 7 and 12) with some localizedrounding of themargins. Even very close to the paleo-seafloor angulargeometry prevails (Fig. 11C; Hurst et al., 2006). This relationshipappears to hold independent of mudstone composition, for example,Miocene biosiliceous mudstone of the Santa Cruz Mudstone Member(Boehm and Moore, 2002; Hurst et al., 2006), clay-mineral-rich lateCretaceous to early Paleocene mudstone with varying biosiliceouscontent in the Moreno Formation (McGuire, 1988; Bartow, 1991) andthe clay-mineral-rich Upper Jurassic Hareelv Formation (Surlyk et al.,2007) all have similar textural features. Quantifying erosion ratesduring sand injection into mud/mudstone at different flow velocitiesand for various host strata properties may be useful experiments todesign as an aid to understanding the formation of natural erosivefeatures.

In our review of outcrop data a relationship between architec-tural elements and the internal sedimentary structures in sandstoneintrusions is encountered that demonstrates a coupled relationshipbetween sedimentation processes and the type of intrusive body. Animportant observation is the presence of laminae in sills (Oberme-ier, 1996; Kawakami and Kawamura, 2002; Diggs, 2007). Laminaeare attributed to pulses of injection (Diggs, 2007), shearing of agranular fabric (Archer, 1984), and variations in both velocity andviscosity of the injecting sand (Peterson, 1968; Taylor, 1982).During deposition the origin of laminae can be explained by upper-stage plane bedform traction (Allen and Leeder, 1980) or shearingof liquefied flows (Lowe, 1976) both of which may occur during thefluidization and injection of sand. Laminae formed by traction sensustricto involve the settling of grains out of suspension onto fluid-sediment interfaces and their movement as upper-stage planebedforms. This mechanism assumes that the flow concentration andthe suspended-load fallout rate are relatively low as not to overloadthe bedload layer by dampening of near-bed turbulence (Lowe,1988). Conversely laminae may form by shearing of concentratedliquefied sand. The sediment behaves as a viscous fluid thatundergoes hydroplastic laminar flow under an applied force(Lowe, 1976; Allen, 1984). Laminae, defined by alignment of platyparticles (mica, some organic matter, mudstone clasts) are observedclose to the margins of liquefied flow-deposits and are attributed tohydroplastic shear induced by frictional interaction between theliquefied flow and its margins (Jones and Rust, 1983).

Because laminae can be generated by flows with very differentrheological characteristics and driven by different processes, deter-mining the mechanism is important as the interpretations imply verydifferent flow regimes. Experiments help to constrain the origin of thelaminae but mainly have been performed in flume tanks and havefocused on how grain fabrics of laminae correlate with the dynamicbehavior of upper-stage plane bedforms (Yagishita et al., 2004). Giventhat the boundary conditions in these experiments don't accuratelycapture those during sand injection there is a need for research thatexamines the hydrodynamics of fluidized sand-flow in fractures.Experiments should focus on developing stability fields for sedimenttransport modes under different flow conditions, such as varying

velocity, grain concentration, viscosity, and fluid pressure, in fracturesof variable roughness, aperture, and orientation.

Few experiments have attempted to model hydraulic fracture andinjection of fluidized sand into low-permeability strata (Nichols et al.,1994; Rodrigues et al., 2009). Rodrigues et al. (2009) injectedfluidized sand and glass microspheres (0-160 μm grain diameter)into hydrofractures (N1 mm wide apertures). These apertures are toosmall to test bedform stability and future work could focus on thenature of sediment transport in fractures of larger aperture (N0.1 mwide).

6.4. Sand extrusion

Sand extrusion has a long-established relationship with seismic-ity (Obermeier, 1996; Galli, 2000) and the occurrence of sandvolcanoes (sand blows) has a variety of applications for predictingthe strength of seismicity and the depth to seismic epicenters(Obermeier, 1996, 1998; Rosseti, 1999; Obermeier et al., 2005). Therelationship between sand extrusion and underlying sand injectiteswas proved by detailed studies that included excavation (Obermeier,1996, 1998; Obermeier et al., 2005) however these studies indicatethat the sand involved in sand extrusion is of shallow (1 to 10's m)origin. This contrasts with the relationships documented at outcrop(Vigorito et al., 2008; Vigorito and Hurst, 2010) and in subsurfacestudies (Huuse et al., 2005) where sand is known to have beentransported through predominantly fine-grained overburden for inexcess of 1 km. It appears that similar sand extrusions can have verydifferent origins, either deep or shallow. Ultimately the depth fromwhich the sand and fluids involved are derived is a function of thedepth at which regional hydrofracture of the seal lithology occurs(Vigorito and Hurst, 2010). In the light of these observations it maybe that the models used to relate the occurrence of sand extrusionsto estimation of seismic magnitude and location require reexamination.

The earliest observation of sand volcanoes in the geological recordare associated with sudden loading caused by the mass-wasting of asubmarine slope (the Carboniferous of County Clare, Ireland; Gill andKuenen, 1957; Strachan, 2002; Jonk et al., 2007). Sub-aerial masswasting of aeolian dunes along an ancient desert margin caused theextrusion of a single discrete sand unit, which is believed to have beentriggered by increased humidity and inherent gravitational instabilityand unrelated to seismicity (Glennie and Hurst, 2007; Hurst andGlennie, 2008). In both examples seismicity cannot be inferred orassumed to have triggered the mass wasting although if thesedimentary loadingwas of sufficient magnitude and occurred rapidlyit may have created a localized seismic event. In contrast the stackedsandstone extrusions near Santa Cruz are probably related to periodsof seismicity along the San Andreas and San Gregorio fault zones(Boehm and Moore, 2002; Hurst et al., 2006) although preciserelationships between ancient seismic events and sand extrusioncannot be established. Seismicity is considered likely to have causedsand fluidization in eolian sandstones from the Navajo Formationalthough precise relationships are not defined (Bryant and Miall,2010). The Danian PGIC has no known affinity with seismicity but hasbeen related to a regional over-pressuring event caused by theeastward subduction of the Pacific plate along the paleo-Californiacoast (Vigorito and Hurst, 2010).

Sandstone extrudites are significant stratigraphic time-lines insedimentary basins as in a geological context they form instanta-neously (Hurst et al., 2006). As elevated, sometimes supralithostatic,pore pressure is the fundamental cause of sand fluidization andinjection (Duranti et al., 2002a,b; Jolly and Lonergan, 2002; Duranti,2007), and this can occur regionally (Huuse et al., 2005; Vigorito andHurst, 2010), crustal processes that create sudden rises in porepressure are the determining factor behind sand intrusion andextrusion. Seismicity may play a role in this but, in particular, when


considering regionally-developed sand injectites other factors arelikely to contribute. The sparse documentation of sand extrudites inthe rock record (Appendix E) severely limits the value of sandstoneextrusions when interpreting shallow crustal evolution and limits ourability to identifying them in the rock record.

6.5. Implication of sand injectites on post-sand-injection fluid-flow

Sand injectites focus fluid-flow through thick sequences (N0.7 kmin North Sea examples, Huuse et al., 2005 and N1.2 km in the PanocheGiant Injectite Complex (PGIC), Vigorito et al., 2008) of fine-grainedhost strata and can be demonstrated on reflection seismic data (Hurstet al., 2003b; Løseth et al., 2003), at outcrop (Jenkins, 1930; Boehmand Moore, 2002; Schwartz et al., 2003) and from petrographic data(Jonk et al., 2003, 2005a,b). This evidence supports the role of sandinjectites in compromising the seal quality of fine-grained strata(Hurst et al., 2003b; Cartwright et al., 2007). Duration of post-injection leakage of fluid (including methane) through the PGIC isestimated as nearly 2 Ma (Minisini and Schwartz, 2002). A sandstonedike intruded into Upper Jurassic mudstones that have not beingheated tomore than ~50 °C contains fluid inclusions that formed fromhotter (~70°–100 °C) basinal fluids thus recording that the dike actedas a conduit for upward flow of fluids with a deeper origin (Jonk et al.,2003). The highly connected, permeable networks created in fine-grained host strata by sandstone intrusions cause muds/mudstones todrain laterally and more rapidly than by compaction-driven verticaldrainage. The possible influence of sand injectites on the thermalconductivity of fine-grained strata is unquantified.

In the context of hydrocarbon exploration the role that sandinjectites, and more specifically sandstone intrusions, play inhydrocarbon migration is obvious but has received little attention(Jenkins, 1930; Hurst et al., 2003b). Sandstone intrusions form highly-permeable (core plug samples average N2D, Duranti et al., 2002a;Briedis et al., 2007; outcrop samples average ~400 mD, Scott et al.,submitted for publication) networks through otherwise lower-permeability (often b0.001 mD horizontal and vertical permeabilitystrata. The five to six orders of magnitude difference in permeabilityhas a huge influence on rates of fluid drainage through fine-grainedstrata that will greatly enhance rates of hydrocarbon migrationthrough what otherwise may be considered as seals or aquitards.

Not only do sand injectites compromise seal quality they can formshallower-than-expected hydrocarbon traps (intrusive traps, Hurst etal., 2005; Frey-Martínez et al., 2007) when intrusive complexes and/or extrudites are sealed by overlying fine-grained strata (Fig. 11C,D).Drilling into reservoir targets that are modified by sand injectiontypically encounters hydrocarbon-saturated sandstones in the fine-grained section (seal!) directly above it. Quantification of the volumeof sandstone associated with sandstone intrusions located abovelarger sandstone bodies in the subsurface, such as those recognizableon reflection seismic data, is challenging as small and steep sandstoneintrusions (for comparison see Figs. 2A and 13) are not resolved onseismic data (Huuse et al., 2007) and frequently not cored as theyoccur higher in the stratigraphy than the levels at which hydrocarbonsare prognosed to occur. Outcrop data confirm the occurrence ofcomplex, often very small aperture (b0.25 m) intrusions (e.g.Figs. 10A and 13B).

Apart from noting the significance of carbonate cement as a controlof the fluid-flow capacity of sand injectites (Jonk et al., 2005a) othergranular-scale characteristics (i.e. grain-size, packing, sorting, andtextures) developed during sand injection are largely ignored in thesand injectite literature (Duranti et al., 2002a,b; Diggs, 2007).Granular fabric in sandstone intrusions have sedimentary structuresthat are formed by variations in grain packing and sorting (Scott et al.,2009), which create permeability heterogeneity at sub-0.1 m scale.Where pipes formed by post-sand-injection elutriation occur en-hancement of sub-0.1 m scale vertical permeability (Hurst and Buller,

1984) is anticipated. The relationship between granular textures andthe permeability of sandstone intrusions is investigated elsewhere(Scott et al., in review).

7. Conclusions

Sand injectites are far more common in the geological record thanthe sparse documentation they have been afforded until the lastdecade. Their increasingly common identification in the subsurfaceassociated with hydrocarbon exploration and production has spurredrecent interest. Sandstone dikes have a long history of occurrence butsills and more irregular sandstone bodies are more recentlyrecognized; sills may have been misinterpreted as depositionalsandstones due to their bedding concordance and internal structuresthat are superficially similar to depositional structures. Recognizingsand injectites and their significance in basin evolution remainschallenging.

Hydrofracture of a seal lithology promotes the upward injection ofsand into an open fracture network. For this to happen, pore pressureexceeded the fracture gradient and sand fluidized such that grainsbecome suspended in fluid by drag forces imparted by upward fluid-flow. In some cases supralithostatic pressure occurs in which case σ1,σ2 and σ3 are approximately equal, the pore pressure can lift theoverburden, sills form and fractures and resultant dikes tend to beirregular and randomly oriented; in the Panoche Giant InjectiteComplex the interval in which this occurs is termed the sill zone.Many sandstone intrusions, in particular dikes that are steeply-inclined to bedding, are planar and orient approximately parallel andperpendicular to the prevailing horizontal stress-axes howeveralmost all sandstone intrusions modify the contact area with theirhost strata, in some instances forming spectacular erosive features.Some dikes and other intrusive bodies are highly irregular andbulbous and lack any obvious relationship to planar, fracture-relatedgeometry; the conditions of their emplacement remain uncertain.

Erosive contacts between sandstone intrusions and their typicallyfine-grained host strata preserve features that formed by forceful andprobably high-velocity fluidized flow during sand injection. Manyfeatures are similar to those formed by depositional processes andmay have formed analogously. The lack of relevant quantitative data,including laboratory experiments, limits interpretation. Evidence of aturbulent flow regime during sand injection is prevalent.

Recent descriptions of the Panoche Giant Injectite Complex (PGIC)allow a better understanding of sand injectite architecture andgenesis and provide a basis for investigating other less well-exposedcomplexes. The tripartite architecture – parent sandstone units,intrusive complex and sand extrusions – described from PGIC is likelyto occur in other less well-exposed, regionally-developed sandinjectites although sand extrusions are not always present. There isa clear imbalance between descriptions of sand injectites from thesubsurface, in particular from reflection seismic data, and outcropsand very few outcrops approach the scale known to exist in thesubsurface.

Relatively little is known about the microscale character of sandinjectites. This limits the understanding of their role in basin evolutionincluding if and how they influence periods of large-scale fluidmigration. In general it seems that little or no attention has been givento their genesis and significance in basin evolution.

Acknowledgements

Much of this work was conducted as part of the Sand Injectitesphase 2 at the University of Aberdeen, funded by DECC (formerlythe UK Department of Trade and Industry), DONG, Lundin,Marathon, Statoil and Total. Review and editorial comments by ananonymous reviewer, Brian Pratt and Andrew Miall are gratefullyacknowledged.


Appendices: synthesis of the literature review

Appendix ALocation, host-strata age, geometry, and internal sedimentary structures of dikes from outcrop and subsurface data.

Author (s) Location Age Geometry Injectite margins and internal sedimentary structures

Newsome (1903) Santa Cruz,California.

Upper Miocene. Bifurcation and rejoining. Thickness between0.6 and 2 m.

Angular shale clasts.

Peterson (1968) SacramentoValley, California.

Upper Cretaceousto Paleocene.

Bifurcation. Thickness between 0.01 and 3 m.Length between 0.5 and 5000 m.

Fluted and grooved margins, alignment of tabulargrains, layering, and grading.

Smyers andPeterson (1971)

Panoche hills,California.

Upper Cretaceousto Paleocene.

Bifurcation. Thickness between b0.1 and 7 m.Length between 1 and 800 m.

Structureless. Ripple marks near contact and angularto rounded shale clasts.

Truswell (1972) Coffee bay, SouthAfrica.

Permian. Ptygmatic folding and boudinage. Predominantly structureless.

Lewis (1973) Oamaru, NewZealand.

Oligocene. Sharp, planar, and irregular. Thickness between0.001 and 10 m. Up to 10 m length.

Lobate and fluted margins. Layering with diffuse andsharp boundaries. Layering is both perpendicular andparallel to dike margins.

Hiscott (1979) Quebec, Canada. Ordovician. Ptygmatic folding and tapering. Up to 2 m length. Structureless.Hannum (1980) Kane County,

Utah.Jurassic. Bifurcation and tapering. Thickness of 0.3 m.

Up to 30 m length.Alignment of tabular grains, grading (finer towardsmargins), and shale clasts.

Taylor (1982) Alexander Island,Antarctica.

Upper Jurassic. Sinuous and planar. Bifurcation. Thickness between0.1 and 5.5 m. Length between 15 and 305 m.

Fluted and grooved margin, alignment of tabular grains,layering, grading, vertical laminae, and angular shale clasts.

Archer (1984) Rosroe, WesternIreland.

LowerOrdovician.

Up to 0.3 m thick. Laterally stepped. Banding parallel to margins and angular mudstonebreccias.

Martill andHudson (1989)

Dossthorpe,England.

Jurassic. Up to 0.2 m thick. Up to 500 m length.Wedge-shaped.

Lobate and slickensided margins. Shale clasts.

Dixon et al. (1995) Bruce-Beryl, UKCS. NA Planar, en-echelon, and tapering. Subangular to angular shale clasts. Conjugate shear fractures.Obermeier (1996) Charleston, South

Carolina.Holocene. Oval- to tube-shaped. Normal grading and organic-rich clasts.

Obermeier (1996) New Madrid,Missouri.

Holocene. Thickness between 0.001 and N100 m wide.Up to 6 m length. Bifurcation and tapering.

Alignment of tabular grains, grading, margin parallellaminae, and pipes.

Tanner (1998) Islay, WestScotland.

Late Proterozoic. Thickness between b0.001 and 0.01 m. Lengthbetween 0.001 and 0.006 m. V-shaped to bulbouswith tails 0.001 m long. Upward pointing cusps.

Structureless. Alignment of tabular grains, smearedshale clasts, angular brecciated shale clasts, microsand-filled fractures, and compositional variations.

Rosseti (1999) Northern Brazil Cretaceous. Up to 0.4 m thick. Up to 0.6 m length. Steeplyinclined and irregular.

Structureless. Convoluted laminae.

Moretti (2000) Brindisi, SouthernItaly

Late Pleistocene. Thickness of 0.1 to 0.5 m. Length between0.1 and 2 m. Rapid thickness changes.

Structureless.

Rodriguez-Pascuaet al. (2000)

Southeast Spain. Late Miocene. Thickness between 0.02 and 0.21 m. Lengthbetween 0.3 and 2 m. Planar to sinuous.

Structureless.

Surlyk and Noe-Nygaard (2001)

Jamesonland, EastGreenland.

Late Jurassic. Thickness of 0.0001 and 1 m scale. Bifurcationand ptygmatic folding.

Structureless. Fluted, grooved, and frondescent markedmargin. Angular, platy, and deformed shale clasts.

Hillier andCosgrove(2002)

Alba Field, UKCS. Eocene. Thickness between 0.002 and 30 m thick.Length between 0.2 and 400 m.Planar and curvilinear margins.

Structureless. Fluted and gutter marked margin.Alignment of tabular grains. Angular shale clastsand flow banding.

Kano (2002) Kukedo, Shimane,Japan.

Miocene. Thickness 0.5 and 1.5 m. Up to 20 m length.Bifurcation

Structureless. Grading and brecciated andesitic clasts.

Strachan (2002) County Clare,Ireland.

Carboniferous. Thickness 0.15 and 1 m. Length between0.4 and 1.5 m. Planar.

Structureless.

Purvis et al.(2002)

Gryphon Field,UKCS.

Upper Paleoceneto Eocene.

Planar and sharp margins. Angular shale clasts with wispy and fine margins.

Jonk et al. (2003) Helmsdale,Scotland.

Upper Jurassic. Thickness between 0.1 and 2 m. Up to200 m length. Bifurcation.

Structureless. Banding parallel to margin.

Mazzini et al.(2003a,b)

Gryphon Field,North Sea.


Thickness between 0.01 and 1 m.Sharp indented margins.

Structureless. Deformation bands.

Duranti and Hurst(2004)

Alba Field, UKCS Late Eocene. Thickness between 0.001 and N2 m.Tabular shape.

Angular and platy shale clasts. Deformation bands andcrosscutting pseudo laminae.

Jonk et al. (2005b) South VikingGraben, UKCS.

Lower Eocene. Ptygmatic folding. Structureless. Angular mudstone clast breccias.Brittle fracturing.

Ross and White(2005)

Coombs and AllanHills, Antarctica.

Jurassic. Thickness between 2 and N5 m.Length between 20 and 75 m. Bifurcation.

Structureless. Grading, banding of differing materialsand margin parallel laminae.

Neuwerth et al.(2006)

Cauca Valley,Colombia.

Plio-Pleistocene. Thickness between 0.05 and 0.3 m. Lengthbetween 0.1 and 1 m.

Structureless.

Briedis et al.(2007)

Balder Field,UKCS.

Eocene. Up to 0.01 m thick. Alignment of tabular grains, banding, and angular toplaty shale clasts. Brittle fractures.

Diggs (2007) Marathon basin,Texas.

Carboniferous. Thickness between 0.01 and 2 m.Planar to sinuous.

Rill marked margin. Alignment of tabular grains.Grading, margin parallel laminae, pipes, and clay clasts.

Glennie and Hurst(2007)

Hopeman,Scotland.

Late Permian. Thickness between b0.005 and 0.4 m.Planar with tapering.

Structureless.

Gożdzik and VanLoom (2007)

Bełchatów mine,central Poland.

Pleistocene. Up to 10 m width. Up to 20 m length. Structureless. Banding parallel to margins.

Hildebrandt andEgenhoff (2007)

Cienega andVitichi, Bolivia.

Middle to UpperOrdovician.

Thickness between 0.3 and 0.5 m.Length between 0.5 and 1.5 m.Sharp contacts with tapering.

Angular claystone clasts.

Hubbard et al.(2007)

Magallanes Basin,Chile.

Cretaceous. Thickness between 2 and 67.Up to N100 m length. Bifurcation.

Fluted and grooved margin, grading, alignment oftabular grains, banding, parallel laminae, and shale clasts.

Jonk et al. (2007) Ross, CountyClare, Ireland.

Carboniferous. Thickness between 0.05 and 0.2 m. Structureless.

Appendix A (continued)


Lonergan et al.(2007)

Gryphon Field,UKCS.


Thickness between 0.1 and 8 m.Ptygmatic folding.

Inclined laminae and angular shale clasts.

Macdonald andFlecker (2007)

Schmidt,Sakhalin.

Miocene. Thickness between 1 and 10 m.Anastomosing.

Eroded margins, grading (reverse and normal) ofangular shale clasts, and margin parallel laminae.

Parize et al.(2007b)

Southeast France. Cretaceous. Thickness between 0.001 and 1 m.Ptygmatic folding.

Structureless. Shale clasts found at junction betweena dike and sill.

Satur and Hurst(2007)

Sleipner Øst Field,Norway.

Paleocene. Ptygmatic folding and irregular margins. Structureless. Inclined laminae, pipes, and angularshale clasts that display jigsaw configurations.

Appendix BLocation, host-strata age, geometry, and internal sedimentary structures of sandstone sills from outcrop and subsurface data.


Truswell (1972) Coffee Bay, SouthAfrica.

Permian. Thickness between b0.5 and 3 m.Bifurcation.

Structureless. Shale clasts observed at the upper contactand at points of bifurcation.

Hiscott (1979) Cap Ste-Anne, Quebec,Canada.

Ordovician. Up to 3 m thick. Stepped andpinch-and-swell thickness changes.

Structureless. Sharp flat margins. Sheet structures.

Archer (1984) Rosroe PeninsulaWestern Ireland.

Ordovician. Up to 0.4 m thick. Stepped, tapered,and splitting. Sharp margins.

Erosional upper margins. Alignment of tabular grains,bands, and shale clasts that display a jigsaw configuration.

Obermeier (1996) New Madrid, Missouri. Holocene. Up to 0.7 m thick. Length betweenb1 and 25 m. Stepped margins.

Erosional upper margins, laminae of lignite and silt/finesand, and vertical gradation of clay clasts.

Surlyk andNoe-Nygaard(2001)

Jamesonland, EastGreenland.

Late Jurassic. Thickness between 0.001 and 1 m.Curvi-linear margins.

Consolidation laminae, dishes, and angular to tabularshale clasts.

Kawakami andKawamura (2002)

Honshu Island,Northeast Japan.

Triassic. Thickness between 0.01 and 0.7 m.Sheet- to wedge-shaped.

Erosional margins. Alignment of tabular grains. Folded,imbricated shale clasts and parallel mud laminae.

Hillier and Cosgrove(2002)

Alba Field, UKCS. Eocene. Thickness between 0.005 and 0.3 m.Bifurcation and tapered tips.

Structureless. Angular shale clasts with a jigsawconfiguration.


Alba Field, UKCS. Late Eocene. Thickness between 0.001 and N2 m. Structureless. Deformation bands. Angular and tabularshale clasts.

Neuwerth et al.(2006)

Cauca Valley, Colombia. Plio-Pleistocene. Thickness between 0.02 and 0.5 m.Length between 0.01 and 5 m.

Structureless. Deformed laminae.

Briedis et al. (2007) Balder Field, UKCS. Eocene. Thickness between 0.01 and 11 m thick. Alignment of tabular grains. Banding, angular andplaty shale clasts with brittle fractures in some clasts.

Diggs (2007) Marathon basin, Texas. Carboniferous. Thickness between 0.001 and 2.5 m.Bifurcation.

Structureless. Alignment of tabular grains. Laminaeparallel to margin.

Macdonald andFlecker (2007)

Schmidt, Sakhalin. Miocene. Stepped. Erosional upper margins.Abrupt thickness changes.

Tool marks and grooves at the margin. Normal to inversegrading of shale clasts and crude parallel laminae.


Gryphon Field, UKCS. Upper Paleoceneto Eocene.

Thickness between 0.1 and 8 m. Undulose laminae and angular shale clasts.

Parize et al.(2007a)

Southeast France. Cretaceous. Thickness between 0.1 and 12 m.Bifurcation, tapered, and stepped.

Tabular with straight margins. Pinch-and-swellthickness changes.

Appendix CLocation, host-strata age, geometry, and internal sedimentary structures of columnar sandstone intrusions from outcrop.


Hannum(1980)

Kane County, Utah. Jurassic. Circular in plan view. Pipe diameter exceeds 1 m. Pipe fill comprises homogenized sandstones and centimeter-to metre-scale blocks of host strata.

Mount(1993)

Adelaide, SouthAustralia.

Lower Cambrian. Circular in plan view. Pipe diameter ranges from 1to 6 cm. Base grades into the surrounding rock.

Concentric internal structure. Central core of very fine sandand a cylindrical to irregular outer halo devoid of fines.

Netoff(2002)

South central Utah. Middle Jurassic. Circular in plan view. Pipe diameter of up to 60 m.Cylinders are near-vertical, pillar-like structures.

Irregular and deformed subhorizontal bedding and lamination(5–15 cm wide). Abundant breccian blocks of host strata.

Huuse et al.(2005)

South East Utah. Middle Jurassic. Circular to oval in plan view. Pipe diametersexceed tens of metre. They cross N100 m hoststrata.

Sandstone beds locally bend up along the pipe margins.Pipe fill comprises homogenized sandstones and centimeter-to meter-scale blocks of dune sandstone.

Chan et al.(2007)

Colorado Plateau. Lower to MiddleJurassic.

Circular in plan view. Pipe diameter rangesfrom cm.size to N10 m. Base flares into thesurrounding rock.

Pipe margins have sharp boundaries. Deformed, irregular,near-horizontal bedding and lamination. Drag structuresfound close to the margins.

Appendix DLocation, host-strata age, geometry, and internal sedimentary structures of irregular injectite geometries from outcrop and subsurface data.


Dixon et al.(1995)

Bruce-Beryl,UKCS.

??? Rounded and discordant margins. Parallel mud laminae. Folded and faulted mudstone raft.Subrounded to angular shale clasts.

Thompson et al.(1999)

Santa Cruz,California.

UpperMiocene.

Up to 24 m thick. Up to 305 m length. Occasional grain-size segregation, banding, laminae,pipes and pillars, angular mudstone rafts, and clasts.

(continued on next page)


Appendix ELocation, age, geometry, and internal sedimentary structures of extrudites from outcrop and subsurface data.

Author (s) Location Age Geometry Margins and internal sedimentary structures


Holocene. Coned. Thickness between 0.3 and 0.7 m.Extends laterally for 5 to 20 m.

Normal gradation of coarse claystone clasts from feederdike. Planar to wavy laminae dipping away from vent.

Takahama et al.(2000)

Karasu Riverdistrict, Japan.

LateQuaternary.

Mounded to ring type structures. Vertical alignment of tabular grains. Angular soil andpumice clasts.

Boehm and Moore(2002)

Santa Cruz,California.

UpperMiocene.

Up to 3 m thick. Extends laterally up to 100 m.Sheet structure.

Parallel laminae and cross bedding. Bioturbation.


Cienega and Vitichi,Bolivia.

Ordovician. Thicknesses between 0.1 and 1.2 m. Basal marginsare downward convex. Upper margins are sharpand flat. Extends laterally 100 to 2000 m.

Angular and platy claystone clasts. Laminated sandy clastsaligned in distinct layers.

Hurst (2004) Santa Cruz,California.

UpperMiocene.

Irregular lower basal surface. Sheet structure. Parallel laminae and cross bedding. Rounded shale clasts.Mudstone breccias. Bioturbation.

Hurst et al. (2006) Santa Cruz,California.

Miocene. Up to 5 m thick. Shallow inclined upper surface.Irregular basal surface.

Low angle laminae and cross bedding. Bioturbation.Shale clasts and breccias.

Glennie and Hurst(2007)

Hopeman, Scotland. Late Permian. Thickness between 0.2 and 0.3 m. Cone-shaped. Low angle laminae (b15˚) that dips away froma central vent. Angular clasts of laminated sandstone.

Jonk et al. (2007) County Clare,Ireland.

Carboniferous. Domed. Up to 0.4 m thick. Up to 1.5 m diameter. Cratered centre with runnels on the dome's slope.

Jonk et al. (2007) Freagh, CountyClare, Ireland.

Carboniferous. Domed. Up to 0.2 m thick.Diameter between 0.3 and 0.4 m.

Undulose laminae.

Jonk et al. (2007) County Clare,Ireland.

Carboniferous. Domed. Thickness between 0.1 and 0.45 m.Diameter between 0.3 and 1.8 m.

Cratered centre with runnels on the dome's slope.

Pringle et al.(2007)

County Clare,Ireland.

Carboniferous. Coned. Up to 0.75 m thick. Up to 3 m diameter. Inclined shale and sand laminae. Cratered centre withrunnels on the domes' slope.

Appendix D (continued)


Surlyk andNoe-Nygaard (2001)

Jamesonland,East Greenland.

Late Jurassic. Rounded margins. Occasional concaveupward margins.

Structureless. Occasional convoluted and consolidationlaminae and dishes. Angular shale clasts.

Diggs (2007) MarathonBasin, Texas.

Carboniferous. Rounded and irregular margins.Thickness between 0.35 and 1.5 m.

Structureless. Occasional faint laminae.Alignment of tabular grains.

Hamberg et al.(2007)

Cecile Field,Denmark.

UpperPaleocene.

Mounded erosional upper margin. Irregular lowermargin. Up to 600 m diameter. Up to 77 m thick.

Parallel and folded muddy laminae. Banding. Angular andplaty shale clasts with sandstone-filled microfractures.


Gryphon Field,UKCS.

UpperPaleocene toEocene.

Irregular and rounded margins. Structureless. Brecciated mudstones. Angular shale clasts.

Appendix FLocation, age, depositional element, geometry, and internal sedimentary structures of remobilized parent units from outcrop and subsurface data.

Author (s) Location Age Depositionalelement

Geometry Margins and internal sedimentary structures

Dixon et al. (1995) Bruce-Beryl, UKCS. Early Eocene. Tabular turbidite. Mounded top. Inclined laminae, dishes, conjugate-shear-fractures, and deformation bands.

Obermeier (1996) Charleston, SouthCarolina.

Holocene. Barrier-bar. Irregular rounded top. Structureless sands.

Surlyk and Noe-Nygaard(2001)

Jamesonland,Greenland.

Late Jurassic. Channel. Mounded top andstepped margins.

Loaded bases, convoluted laminae,consolidation laminae, dishes, pillars,and shale clasts.

Hillier and Cosgrove(2002)

Alba Field, UKCS. Late Eocene. Channel. Mounded top and winged. Dishes and pillars.

Purvis et al. (2002) Gryphon Field, UKCS. Upper Paleoceneto Eocene.

Turbidites andsandy debrites.

Mounded and steep-sided. Inclined clay-enriched laminae and cusps.

Strachan (2002) County Clare, Ireland. Carboniferous. Tabular turbidites. Domed top. Folded bedding, normal and reverse faulting,and boudinage.


Alba Field, UKCS. Late Eocene. Channel. Mounded top and winged. Consolidation laminae, inclined laminae,dishes, pillars, and mudstone clast breccias.

Briedis et al. (2007) Balder Field, UKCS. Upper Paleoceneto Eocene

Gravity-flowdeposits.

Mounded top and steepmargins.

Structureless sandstone and pipes.

Hamberg et al. (2007) Cecile Field, Denmark. Upper Paleocene. Channel. Mounded top and steepmargins.

Structureless sandstones with minor faintbedding.


Cienega and Vitichi,Bolivia.

Middle to UpperOrdovician.

Shoreface. Loaded top. Structureless sandstones and minor faintbedding.

Hubbard et al. (2007) Puerto Natales, SouthChile.

Cretaceous. Channels. Steep margins and winged. Structureless conglomerate.

Lonergan et al. (2007) Gryphon Field, UKCS. Upper Paleoceneto Eocene.

Sandy debrites. Steep discordant top. Structureless sandstones, pillars, pipes,and shale clasts.

Satur and Hurst (2007) Sleipner Øst Field,Norway.

Paleocene Gravity-flowdeposits.

Rapid thickness changes. Consolidation laminae and dishes.

Surlyk et al. (2007) Jamesonland,Greenland.

Late Jurassic. Moundedsand-bodies.

NA NA

Surlyk et al. (2007) Jamesonland,Greenland.

Late Jurassic. Lateral extensivesand-bodies.

NA NA


Appendix GEvidence of flow regime during sand injection from outcrop and subsurface data.

Author (s) Location Element(s)

Nature of flow Evidence

Dott (1966) – Dike. Laminar. Internal lamination and grain alignment.Peterson (1968) Sacramento

Valley,California.

Dike. Laminar. Grain alignment, layering, and graded layering.

Taylor (1982) AlexanderIsland,Antarctica.

Dike. Laminar. Smooth dike walls, laminae sub-parallel to parallel to the dike walls,and coarser-grained fraction in the central parts of the dikes.

Dixon et al. (1995) Bruce-Beryl,UKCS.

Dike andsill.

Laminar. Concentration of coarser host shale clasts in the central portions of the dike.Alignment of tabular shale clasts and lamination.


Dike,extrudite.

Initially turbulent. Fining-upward sequence from basal clay clast-bearing sand within the dikeand lower extrudite to clean medium-grained structureless sands.

Kawakami andKawamura (2002)

Honshu Island,Northeast Japan.

Sill. Turbulent. Upper flute-like erosional margins, rip-up clasts of host mudstone, and mudclast rich detrital matrix.

Diggs (2007) Marathon basin,Texas.

Dike. Laminar flow. Alignment of elongate quartz grains, clay chips, and micas parallel to the wallsof the injection. Pore-filling clay matrix and mud clasts.

Hamberg et al.(2007)

Cecile Field,Denmark.

Mound Turbulent flow. Vortex-likeor irregular flow patterns.

Mounded erosional upper margin with ripping and spalling of clasts. Increased mud clastcontent towards to the erosional margins. Rounded incisions along the edges of clasts.

Hubbard et al.(2007)

Puerto Natales,South Chile.

Dike. Turbulent. Rip-up material of host rock and flutelike grooves. Approximate velocity of 0.70 ms−1

based on 20 cm clast size using the square-root law (Allen, 1985).


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physical characteristics of sand injectites

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