Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks
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SEDIMENTS PRODUCED BY IMPAC
655
Krom the study ofthe distribution of microtektites (see below)
in the Australasian strewn field. Glass and Pizzuto (1994)
derived the following equation:
t
=
0.02 Rir/R)'-*'^^''-
(Eq. 2)
which can be rewritten as a function of the distance alone:
t=
10--"'^"^
-'^="-*
(Eq. 3)
This relation seems to be valid not only for proximal, but
also for distal ejecta, such as tektites and microtektites.
Impact ejecta arc found in Archean and Proterozoic rocks
sequences in Australia and South AfVica. Investigations of
these layers arc still in progress (see chapter by Koeberl, this
volume). However, the study of the K T boundary ejecta
provided the rnost influence for the discussion about the
importance of itnpact events. It is easy to deteet the K T
boundary layer in the field, where is appears as a distinct break
in lithology, with a thin (usually up to 1 or 2cni thick) layer
commonly composed of clay or claystone. A detailed deserip-
tion of the field occurrences of various K T boundary layers
relations was recently given by Smit (1999), and references
therein. Information on a variety of aspects ofthe study ofthe
K T boundary, the K T impact event, and its implications
can be found, for example, in the "Snowbird"' series ol'
conference proceedings (Silver and Sehultz, 1982; Sharpton
and Ward. 1990; Ryder etal., 1996; Koeberl atid MacLeod,
2002).
Erom the sedimentological perspective, K-T boundary
layers in marine and terrestrial sections have a somewhat
different appearance. Terrestrial K-T boundaries (e.g., in the
Western Interior of North America) are characterized by an up
to 3-ctn-thick dual layer, whereas in most marine sections it
dties not display a dual-layered nature. At the Western Interior
locations, the lower layer is mainly kaolinitic with some
smectite; it contain hollow spheres up to
1
mm in diameter.
that are most often also composed of (and filled by) kaolinite,
but at some sites (e.g.. in Wyoming) they are composed of
goyazite (a hydrous alumino-phosphate: an alteration pro-
duct).
Bohor and Glass (1995) give a detailed description ofthe
formation and alteration of these spherules. The upper
(thinner) layer is commonly laminated and consists of smeetite
and kaolinite clay with variable amounts oi' organic matter,
giving it a dark gray to black appearance (Bohor, 1990). Most
ofthe shocked minerals and the siderophile element anomaly
(see below) are restricted to the upper layer; it has been called
the "fireball" layer (e.g.. Bohor, 1990). However, the
composition and internal structure ofthe K/T boundary layer
varies with distance from the source and depositional
environment (e.g., Stnit, 1999). In marine sites in Italy, for
example, a cm-thick clay layer occurs within well-bedded
limestones (Eigure S2l). Besides high contents of the PGEs
and the presence of shocked minerals, evidence for impact in
the various K-T boundary layers includes also the presence of
osmium and chrotniutn with extraterrestrial isotope ratios,
fragments of the original bolide, impact-related diamonds.
high-pressure polymorphs of quartz (eoesite, stishovite), and
impact glass. In the early 1990s, the 200-km-dianieter
Chicxulub struettire in Mexico (Yucatan Peninsula) was
confirtned as an itnpaet structure of K T boundary age. Thus,
the source crater of the K T boundary impact was found.
There are several other cases where impact craters were
identified after careful studies of distal ejecta layers. Eor
example, Gostin etal. (1986) reported on the discovery of an
Figure S21 The Cretaceous-Tfrti.iry bound.iry Liy^r at Frontale,
Umbrid-MiircheseqtJencc, central Italy. The grayish clay layer is about
1.5 cm thick and situated within limestones. Encl-Cretateous rocks dre
marked with a K, and earliest Tertiary rocks are marked wilh a T.
impactoclastic layer within late Precambrian shales of the
590 Ma Hunyeroo Formation in the Adelaide geosvncline.
South Australia. This layer was found to contain abundant
shocked quartz grains, zircons, shattered mineral grains, and
small shatter cones (e.g.. Wallace etal.. 1996). The ejecta were
found in outcrops and drill cores over several hundred
kilotneters in south-central Australia. Around the same time,
Williatns (1986) identified the Aeraman structure in South
Australia as an impact structure, which was rapidly confirmed
as the source crater of the Bunyeroo impact ejecta layer.
Gostin et al. (1989) attd Wallace et al. (1990) detected
enrichments of the PGEs in the ejecta layer; however, post-
fbrniational redistribution had altered the PGE patterns. The
diameter of the deeply eroded Aeraman structure is probably
about 90 km. Impact ejeeta have been found at distances of up
to 450 km from the Aeraman structure (i.e.. about 10 crater
radii),
making this a trtie distal ejecta layer.
Another important group of distal impact ejecta are
tef<tites
andtnicrotektites. They are a group of natiual glasses that have
been known for several centuries by now, which earlier had
been interpreted to represent glassy meteorites or ejeeta from
terrestrial or lunar volcanoes. Mainly due to chemical studies,
it is now commonly accepted that tektites are the product of
melting and quenching of terrestrial rocks during hypervclo-
city impact on the Earth. The chemistry of tektites is in many
respeets identical to the composition of upper crustal material.
Tektites arc currently known to occur in four strewn iields of
Ccnozoic age on the surface of the Earth. Strewn fields can be
defined as geographically extended areas over which tektite
material is found. The four strewn fields are: the North
American. Central European (moldavite). Ivory Coast, and
Australasian strewn fields. Tektites found within each strewn
field have the satne age and similar petrotogical. physical, and
chemical properties.
In addition to the (macroscopic) tektites found on land,
rnicrotektites were found in three of the four strewn fields.
Microtektites oecur in deep-sea cores, within layers that
correspond in stratigraphic age to the radiogenic age of
tektites on land, are generally less than
1
mm in diameter, and
show a somewhat wider variation in chemical composition
656
SEDIMENTS PRODLICFI) BY IMPACT
than tektites on land, but with an average composition that is
very close to that of "normal" tektites. Microtektites have
been very important for defining the extent ofthe strewn Iields.
as well as for constraining the stratigraphie age of tektites, and
to provide evidence regarding the location of possible source
craters. Eor details see Koeberl (1994) and Montanari and
Koeberl (2000).
Tektites are chemically homogeneous, often spherically
symmetric natural glasses, with most being a few centimeters
in size. Those found on land have commonly been classified
into three groups (see, e.g., O'Keefe, 1963): (1) normal or
splash-form tektites; (2) aerodynamieally shaped tektites; and
(3) Muong Nong-type tektites (sometitnes alsocalled layered
tektites). The aerodynamic ablation results from partial re-
melting of glass during atmospheric re-entry after it was
ejected outside the terrestrial atmosphere and solidified
through quenching. The shapes of splash-form tektites include
spheres, droplets, teardrops, dumbbells, etc., or fragments of
such shapes. These are not to be confused with aerodynamical
Ibrtns.
Muong Nong-type tektites were named after the type-
locality in Laos. They are usually considerably larger than
normal tektites and are of chunky, blocky appearance. Muong
Nong-type lektites show a layered strticture with abundant
vesieles. These tektites are less depleted in volatile elements
than the splash-form tekfites.
Microtektites are, by definition, less than
1
mm in si/e. and
show a variety of shapes, ranging from spherical to dumbbell.
disc,
oval, and teardrop, in color they range from colorless and
transparent to yellowish and pale brown. They often contain
bubbles and leehatelierite inclusions. Microtektites occur in the
stratigraphic layers of the deep-sea sediments that eorrespond
in age to the radionietrically determined ages of the tektites
found on land. Thus, they are distal ejecta and represent an
impact marker. For example, in deep sea sediments in the
Indian Ocean and Pacific Ocean in the general vicinity of
Australia and Indochina they have a stratigraphic occurrence
near the Brunhes-Matuyama boundary; this and their
composition indicate that they are indeed part of the
Australasian strewn field-
The geographical distribution of microtektite-bearing cores
defines the extent of the respective strewn fields, as tektite
occurrences on land are much more restricted. Eurthertnore.
microtektites have been found together with melt fragments,
high-pressure phases, and shocked minerals (e.g.. Glass, 1989;
Glass and Wu, 1993) and. therefore, provide confirming
evidence for the association of tektites with an itnpaet evenl.
The variation of the microtektite concentrations in deep-sea
sediments with location increases toward the assumed or
known impact location. Glass and Pizzuto (1994) determined
the geographic variation in abundance of Australasian
microtektites (ranging frotn <l to 3.255 microtektites per
cm^ in the > 125 |.tm size range), and found that the abundances
increased toward Indochina; and deduced a possible source
region in Cantbodia.
Within the past few decades the impaet origin of tektites was
established, raising the question regarding their source craters,
Sinee then, numerous suggestions and educated guesses have
been made regarding the location ofthe possible source craters
for the tektite strewn fields. Relatively reliable links between
craters and tektite strewn fields have been established between
the Bosumtwi (Ghana) and the Ries (Germany) craters, and
the Ivory Coast and the Central European fields, respectively.
Only recently, a platisible link was established between the
newly discovered Chesapeake Bay crater and the North
American strewn field. No large crater with a compatible age
has yet been identified for the Australasian strewn field. "I'he
exact process of tektite formation is still not quite elear,
although from chemical and isotopic studies it is fairly well
established that tektites have formed fYom the immediate
surface layers ofthe target rocks {cf. Koeberl. 1994).
Distal ejecta ("impactoclastic layers") can be used as
markers for impact events in the stratigraphie record. "Impact
markers'" can be deseribed as all chemical, isotopic. and
mineralogical species derived from the encounter of cosmic
bodies (such as eometary nuclei or asteroids) with the Earth, as
discussed in more detail by Montanari and Koeberl (2000).
Such markers are quite important in the detection and study of
aecretionary events in the sedimentary record, to identify their
origin, and to evaluate their possible role in global change and
on the Earth's biotic and climatie evolution throughout
geological time. Distal ejecta layers can be used to study a
possible relationship between biotic changes and impact
events, because it is possible to study such a relationship in
the same outcrops, whereas correlation with radiotnetric ages
of a distant impact structure is always associated with larger
errors.
Impactoclastic layers arc composed of distal ejecta. For
example, microtektites oecur in well-defined and thin layers in
deep-sea sediments and provide an excellent time marker. The
sedimentology of impaetoelastic layers can provide important
informatiott not only on source craters and impact proeesses,
but also on the effects (local, regional, or global) of the related
impact events.
Christian Koebcr!
Bibliography
Alviirc/, L.W., Alvarez, W,. Asaro. F.. imd Midid, H.V., 1980.
Extraterrestrial cause for the Cretaceous-Tertiary extinction.
Science. 208: 1095-1108.
Alviire/, W., 1997.
T.
Re.xaitd
the Crater
of
Doom.
Princcloit University
Press.
Bohor, B.F., 1990. Sliocked quartz iitid more: inipuct sigtiatures in
CrL'taL'eoiis/Ternitry botindary clays. In .Shitrpion. V.L.. and Ward,
P.D.
(cds.).
Global
Catastrophesin Earth History, Geological Society
of America, Special Paper, 247. pp. 335-347.
Bohor. B.F.. and Glass, B.P.. I99.S. Origin and diitiienesis of tlie K/T
impacl sphcrtilcs From Haiti to Wyoming and beyond. Mcteori-
tics.
30: 182-198.
Bohor. B.F., Foord, E.E.. Modrc^ki, P.J.. and Tiiplelioin. D.M,. 1484.
Mineralogieat evidence for an impac! event al the Cretaccotts/
Tertiary boundary. Science. 224: 867-869.
Bohor. B.F., Modre-ski, P.J.. and Foord, E.E., 1987. Shocked qtiartz in
tiic Cretiiceous/Tertiary boundary clays: evidence lor global
distrihntion. .Scietice. 236: 705 70S.
Frencli, B.M.. 1998. Tracesof Catastrophe: A Hamlhookof Shock-Meta-
niorphie ^fleets in Terrestrial Meteorite hnpaet Strnctnrcs. LPI
fontribtnion 954, Hotiston: Ltinar and Planetary Instiltite. 12()pp.
Gliiss, B.P., 1989. Nortli Americiin tektiie debris and impacl ejecta
from DSDP Site 612. Meteoritics. 24: 209 218.
Glass,
B.P.. and Pizznto, J.E., 1994. Geographie variation in
Anslralasiaii mierotektitc concentrations: implications concerning
Ihe location and si/e ofthe source crater. Journal of
Geophysical
Research. 99: 19075-19081.
Glass,
B.P.. and Wu. J-, 1993. Coesite and slioeked quartz discovered
in the Attstralasian und North Atiierican mierotektite layers.
Geology. 2\\ 435-438.
Glen. W., 1994.
The
Mass Extinctitni Debates, Stanford University Press.
Gostin, V.A., Haines. P.W.. Jenkins. R.J.E.. Compston, W.. and
Williams, I.S., 1986. Impact ejecta hori/on within late Preeanihrian
shales, Adelaide Geosyitcline. south Ausiralia. Science. 233:
!98 200.
SFPTARIAN ("ONCRFTIONS 657
Gosiin. V.A.. Kcays. R.R.. and Wallncc. M.W,. 11S9. Iridiiim
anomaly from the AcTaman impact eJLVla horizon: impafls can
produce scdimciiUiry iridkim peaks. Ntiliiir. 340: 542 544.
Grieve. R.A.F., Lanjienhorst. F.. and Suimcr. D.. ]99<r Shock
mctaniorphism in nature and experiment: II, Signilicarice in
geosciencc. McWoriliisuiulPhiiianryScicini', 31: 6 .15.
Koeberl. C. 19'?4. Tektite origin by hypervelocity astcroidal or
cometaiv impaet: target roeks. source craters, and mechanisms.
In Dressier. B.O.. Grieve. R.A.F.. anti Sharpton. V.I.. (eds.). l.iiriic
MvH'oriie Iniptuis ami Piunviary
F-voiiiiinu.
Geological Society of
.'\merica. Special Paper 2'J.l. pp 1.13- 152.
Koeberl. (.'..
21)01.
The sedimentary record of impact events. In
Peucker-nhrenbrink. B.. and Schmitz, B. (eds.). Aari'lioiiofEMni-
Ivrruslriiil MnHcr
Thriiiii^iiiinl
iMvili's Hislory. New York: Kluwer
Aeademic/Plenum Publishers, pp. 333-378.
Koeberl. C. and MacLeod. K. (eds.). 2002.
CuUistropbiv
Eveiilsiimi
Muss E.xiiiifiii'us:
IIII/KHI.S
iiiul Bcyinul. Geological Soeiely of
America. Special Paper. 356. 746 pp.
Melosh. H.J.. 19S'). liiipiia Crciwriiiii.A Giwingic Pnxi's.s. New York:
Oxl'ord University Press.
Monlanari, A., and Koeberl. C. 2000. /inpiicl Sirciiii'riiphy.Thc
Ilii/iciii
Ream!.
Lecture Notes in Eiarih Sciences. Volume ')3. Heidelberg:
Springer Veriag. 3fi4pp.
Oberbeck. V R.. 1975. The role of ballistic erosion and sedimentation
in lunar stratigraphy. Reviews of
Ct'oplivsics
ami Spaa- Plivsics. 13:
337-362.
O'Keefe. J.A. (ed.). l%3. Tckiiivs. Chicago: University of Chicago
Press.
Peueker-Lhrenbriiik. B.. and Schniit/. B. (eds.). 2001. Accrciion iij
E.xiniii'rnwlricil Maiii-r Tlirimglioiil Etirlh's History. New York:
Kluwer Aeademic/Plenuni Publishers.
Powell. J.L.. 1998. Ni^lil Comes lo iln- Ocuiccoiis. New. York:
W.H. Freeman.
Ryder. G.. Fastovsky. D.. and Gartner. S. (eds.). 1996. The Crckic-
I'otis-Teniary Evenl iiml Other Catastrophes in Earlh Hi.slovy.
Geological Society ol' America, Special Paper. 307. 576pp.
Sharpton. V.L.. and Ward. P.D. (eds.). 1990. Giohol Ctittisimphcs in
Earlli Hislorv. Geological Society of America. Special Paper 247.
631pp.
Silver. L.T.. and Schult/. P.M. (eds.). I9S2. (k'oiojiicaiInipiicatiinisof
Impacts of
Lciriie
.Astcniiils and Comets on tin- Earth. Geological
Society of America. Special Paper 190. 52Kpp.
Smit. J.. 1999. The global stratigraphy of the Cretaceous-Tertiary
boundary impact ejeeta. .liimiiii Revieiis of Earth untl Plaiieuirv
Sciemv.'ll: 75 113".
Sttifller. D.. and Langenhorst. F.. 1994. Shock metamorphism of
quartz in nature and experiment: 1. Basic observations and theory.
Mrteoritits.l9:
155 lSl.
Wallace. M.W.. Gostin. V.A.. and Keays. R.R.. 1990. Acraman
impaet ejecla and host shales: evidenee for low-temperature
mobilization of iridium and other platinoids, (n'oloiiv. 18: 132 135.
Wallace. M.W.. Gostin. V.A.. and Keays. R.R.. 1996. Seditiiciilologv
of the Neoproterzoic Acraman impact ejeeta horizon. South
Australia. .iGSO: Journal .Aiisiraiian (jetihi^v and Geophviis. 16:
44.3 451.
W
illiams. G.E., 1986. The Aeraman impact structure: source of ejeeta
in late Precambrian shales. South Australia.
.Seienee.
233: 200 203.
Cross-references
Exiralerresti'ial Materiiil in Sediments
Keattircs Indicating Impact and Shock Metamorphism
SEPTARIAN CONCRETIONS
Septarian structtircs ai'c former cracks, often lilled with cenienl
and iirc most commoniv ibttTid Iti coticrctions hosted iti
tnudrocks. although rare occurrences are known from
siUstones and sandstones. Septarian structures occur in
concretions or coticretionary sheets, which may be chctnically
and mineralogically the same as non-septarian concretions in
Ihe same mudrock. The host concfetions are most cotnmonly
calcite. dolomite or sideritc dotninaled. aitiu)ugh rare occur-
rences have been interpreted from silica concretions. Septarian
concretions are predominantly a pre-Pleistocene phenomenon.
aUhough a potential "proto"" version has been described tVotn
the late Pleistocene (Duck. 1995).
Crack morphology
Septarian structures were initially formed as open fractures,
and are most often concentrated in ihe central regions of
concrctioits and reduce in width and frequency toward the
outer parts ol" the concretion, which may or may not be
cracked. The fractures take a variety of forms, from lenticular
cracks, which are widest toward the center of the concretion.
to straight-sided cracks, to dense networks of hairline cracks.
The orientation of the cracks can be diverse, but always
include cracks which are predominantly stib-vertical to the
bedding plane. Sotiie sub-vertical craeks in sections perpendi-
cular to the bedding plane may appear inclined, which is often
due to the low angle intersection o\' the craek and section.
Cracks can also be sub-horizontal or curved, sometimes
reflecting the shape of the concretion (Figure S22A). Within
horizontal sections the cracks often demarcate approximately
equidirnensional blocks of the concretion (Figure S22B).
Multiple generations of cracks are apparent in some septarian
concretions, cross cutting earlier formed cracks. Cracks are
tnostly displaeivc and may cut body fossils, or occasionally
cracks may show a component of shear displacement.
Crack filling cements
Cracks may range iVom largely unfilled to fully cetnent filled,
often with a variety of distinctively colored spar cements. The
Tills may also include geopetal seditnent. or sediiitcnt injected
frotn the surrounding tnudrocks. through cracks that reached
the concretion surface. I'ltc cements grew on the fracttire walls
within a pore-fiuid tilled craek. This is evident by the
preservation o( partially unfilled cracks displaying crystal
terminations in the voids. The cracks are usually filled with
calcite, but sometimes include other carbonates, sulfide,
sulfates or clay minerals, sometitncs quite different to the
tnineralogy of the concretion body. In most occurrences the
volume of the cracks make up less than 20 percent of
the volume of the concretion body, although in rare
occurrenees the cement-filled cracks may constitute up to
•^70 percent of the concretion body.
Timing of crack formation
The titning of the crack formation relative to the fortnation of
the coneretion body is indicated by the geochemistry of the
body and crack cements. Detailed carbon and oxygen isotopic
studies (see Isolopic Mi'llwds in Si'dtnicntolci^y) indicate the
central region of the coneretion body generally formed first.
The outside of the concretion body may have a similar cement
chemistry to the earher fortncd crack cements, hence for some
concretions the outer part of the concretion formed at the
same time as the inner cracks fills. The fracturing of some of
658
SEPTARIAN CONCRETIONS
(A)
Figure S22 (A) vertical sertion through a Luni|Jt)und p
central parll iind siderite (dark OLiterl seplari.in Luiuretion. The
lenlit uLiranri parallel-sided cemenl-l'illed cracks
are
moslly vertical to
!)eddin}i;.
Three generations of cemenl till occur a clear calcite spar
(dark) in the calcite concretion (large anil small cracks), a [lale brown
sp,ir [pale grayl and a barito (while), which mostly fills the laier
tapering terminations to cracks. Some
L
rat k voids are preserved in the
inner calcite concretion (Ironstone Shales, Lower Jurassic, UK). IB)
horizontal section through dolomilic septarian concretion with
septarian cracks partially (illed by dolomite (Lower Carboniferous oi
Berwickshire, Scotland).
the early cement fills by later cracking, and subsequent filling
by later cements has indicated that some concretions have a
repeated history of cracking and successive crack tills. The
linkage of cement chemistry to likely changes in pore-lltiid
chemistry with sediment burial (see Diagencsi.s). has aiso
indicated that in some septarian concretions the crack-fills
formed later than the concretion body.
Depth of formation
Estitiiates of the depth of formation of septarian craeks are
elosely linked to the origin of the body of [he coneretion. There
are three main cities to the depth of t^orittation. Firstly, the
variation in the amount of cement in the concretion body has
been used as a relative indicator of burial depth, through its
assumed passive filling of the early sediment pore-space. This is
not without its probletiis. due to the possible displacive or
repliicive growth of sotiie concretion body cements. Secondly,
the isotope geochemistry of the crack and body cetnents.
linked to tnodels of organic matter stimulated diagenesis (see
Diuf-cncsis) has suggested estimates of likely burial depth,
based on typieal present-day depths of pore-water chemical
boundaries. Thirdly, occurrences of septarian concretions
exhurned at disconformities and subsequently re-buried, in
which estitnates of the atnouttt of sediment removed can be
ascertained (Hesselbo and Paltner, 1992), These three estimates
suggest that some septarian cracking was initiated at a
niiitimtim of a few meters to a few 10s of tneters depth.
Ma.vitiiuni cracking depths arc ot^ten not well constrained, but
tnay extend to several kms of burial, or occur during uplift.
Origin of cracks
Septarian structures have attraeted debate since the late 19ih
Century, when the accepted hypothesis was that they formed
by shrinkage of the interior parts of the concretion. This idea
has attracted many proponents since that time involving
increasingly more cotnplicated growth models (Pratt, 2(101).
However, a plausible model that explains this shrinkage
process in firm physical terms has largely eluded its propo-
nents.
The physical proeess that may provide a comparable
analogue for the shrinkage hypothesis is the production of
synercsis cracks (see Syuvirsis). The shrinkage mode! is also
deficient in explaining the physieal and chemical nature of the
cementing ttiedium which shrinks in the early-formed parts of
the concretion, and deficieitcies pertaining to the textural and
geochemistry data about the relative timing of textural features
(Astin, 1986). More recent hypotheses suggest the cracks are
generated by lensional failure, either through stresses pro-
duced by excess pore pressure (see Compaction: Asliti. 1986;
Ht)uns[ow, 1997) or shaking motions produced by earthtitiakes
(Pratt. 2001). The essential differences between these ttiodels
are that the Astin and Ptatl tnodels implicate fortnation-wide
stresses, which are through unknown processes, localized to
the concretion body. Without a process of tensional failure
localization, both the Astin and Pratt models would also
produce septarian eraeks in rocks other than the concretion.
The Houiislow model implicates stresses localized to the
cementing concretion, whose cementation is the cause of (he
tensional stress. The driving motivation for tensional fracture
models is the realization that large instantaneous tensional
forces are needed to fracture rigid body tbssils and the
concretion body in its consolidated state. Tensional
stresses applied over a long period of time (sub-critical crack
growth) appear to provide a solution in lowering the
threshold for fraeture growth, and partially cetnented concre-
tions,
as implicated in some growth tnodels. may lower the
tensional strength of the coneretion body, A deficieney in
the tensional fracture tnodels are that in some concretions the
lensional strains implied by the large volumes of crack-filling
cetneitt are far in excess of what could be expected. Hence, a
hybrid shrinkage- tensional fracture model may be appropriate
for some septarian concretions (Pratt. 2001).
Ultimately the origin of septarian cracks is tied to the debate
about the origin of concretion bodies. One of the ptoblems is
the lower nutiiber of case studies linking the timing of
SILCRETE
659
concretion body and septarian cement formation, which utilize
integrated and detailed geochemical, textural and mineral-
ogical data. Study of hiatus septarian concretions could be a
fruitful avenue for further study, since exhumation provides a
relative timing and depth indicator (Hesselbo and Palmer,
1992).
A second gap in knowledge is the physical properties of
concretions when they first form, which is fundamental in
dictating the likely processes of septarian crack formation. The
absence of a present day analogue for septarian concretions
either suggests the formation processes are absent or very rare
at the present day, or that they take longer to form than
current understanding suggests.
Mark W. Hounslow
Bibliography
Astin, T.R., 1986. Septarian crack formation in carbonate concretions
from shales and mudstones. Clays Clay Minerals, 21:
617-631.
Duck, R.W., 1995. Subaqueous shrinkage cracks and early sediment
fabrics preserved in Pleistocene calcareous concretions. Journal
Geological
Society of London, 152: 151-156.
Hesselbo, S.P., and Palmer, T.J., 1992. Reworked early diagenetic
concretions and the bioerosional origin of a regional discontinuity
within British Jurassic marine mudstones. Sedimentologv, 39:
1045-1065.
Hounslow, M.W., 1997. Significance of localized pore pressures to the
genesis of septarian concretions. Sedimentotogy, 44: 1133-1147.
Pratt, B.R., 2001. Septarian eoncretions: internal cracking caused by
syn-sedimentary earthquakes. Sedimentology, 48: 189-213.
Cross-references
Compaction (Consolidation) of Sediments
Diagenesis
Diagenetic Structures
Isotopic Methods in Sedimentology
Syneresis
SILCRETE
Silcretes or siliceous duricrusts are defined by Summerfield
(1983) as the indurate products of surficial and penesurficial
(near-surface) silica accumulation. They are formed by
cementation and/or low-temperature replacement of bedrocks,
weathering deposits, unconsolidated sediments, soil or other
materials. Summerfield proposes silcretes should contain at
least 85 percent silica, however some silcretes may not contain
this percentage as they retain parts of the parent material.
Silcretes form nodular, columnar and/or massive layers in
outcrops. In hand specimens they show many different features
inherited from the parent material or acquired during silica
accumulation and afterward in silica diagenesis. Silcrete
horizons usually vary from several cm to
3
m, but can
occasionally exceed this thickness. Although silcretes exist
globally, the principal studies on them have been carried out in
Australia and Africa.
Once the petrology of a silcrete is known, any further
discussion of its formation needs to explain: (1) the source of
silica; (2) the silica transfer from the source to the site of
accumulation; (3) the conditions and mechanisms that give rise
to silica precipitation; and (4) their relation to the climate and
paleosurfaces.
Mineralogical composition and petrology
The main SiO2 minerals deposited during silcrete formation
are:
quartz, moganite (silica polymorph with a crystal structure
closely related to that of quartz), opal CT (stacked sequences
of cristobalite/tridymite and occasionally Opal A (amorphous
silica).
These silica phases display many fabrics and textures,
depending on the parent material, chemical composition of the
pore fluids, and whether the mineral formation process is
replacement or cementation. Consequently the petrography of
silcretes is highly complex.
Common crystalline quartz of different sizes is frequently
associated with several varieties of chalcedony. Chalcedony is
fibrous microcrystalline quartz and different varieties (with
distinct names in French and English) can be observed using a
polarizing microscope. These are: calcedonite (length-fast
chalcedony, in which the elongation of the fibers is perpendi-
cular to the crystallographic c-axis), quartzine (length-slow
chalcedony, in which the elongation is parallel), lutecite
(another type of length slow, in which the fiber axis is inclined
by approximately 30°) and helicoidal calcedonite or zebraic
chalcedony (systematic helical twisting of the fiber axes around
the crystallographic c-axis). The varieties of chalcedony allow
the definition of formation environments as acid or nonsulfate
(length-fast) or basic or sulfate/magnesium rich-environments
(length-slow) (see review in Hesse, 1991). Unfortunately, there
are exceptions to these rules and rigid application of these
criteria can lead to error in interpretation.
The microscopic characteristics of moganite have not yet
been studied in depth, but it is generally described as uniformly
flaky or fibrous (length-slow). Moganite, in amounts of over 20
percent, is considered indicative of special evaporitic environ-
ments which have still not been clearly determined (Heaney,
1995).
The opal in silcretes is generally opal CT and appears
throughout the rock as massive or glaebular. Glaebules are
concentric rings of silica and matrix components and are
usually less than one centimeter in diameter, although they can
occasionally reach several centimeters. The opaline cements
are fibrous and named lussatite (length-slow) or lussatine
(length-fast). Opal A is not common, but of great interest,
because in some cases its existence propitiates the formation of
opaline gemstones. The opaline phases in silcretes appear
where the parent material is mudstone or the silica concentra-
tion of the pore fluids is very high (120ppm of silica for
amorphous silica and 89ppm for opal CT). These opaline
phases can recrystallize to quartz (aging), tn the case of
silcretes, temperature and time, the two main factors favoring
aging, are not always necessary.
The major silcrete types and genesis
Parent materials determine the majority of the mineralogical
and mieromorphological characteristics of silcretes and as
there are many parent materials, there are many varieties of
silcretes. It is therefore difficult to find a universally acceptable
petrographic classification of silcrete. Summerfield's classifica-
tion (1983) has been commonly used, although this classifica-
tion is not appropriate for silcretes formed from nondetrital
660
SILICEOUS SEDIMENTS
parent materials. Thiry
and
Milnes (1991) developed
a
general
genetic classification
of
silcretes, which subdivided them into
pedogenic
and
groundwater types. Pedogenic silcretes
are the
result
of
silica precipitation within
the
soil profile.
The
structures
and
fabric
of the
parent materials
are
disturbed
or
destroyed during silcrete formation being substituted
by
others
more typical
of
paleosols. Intermittent
and
repetitive leaching,
infiltration
and
illuviation alternate with evaporation, which
causes
the
precipitation
of
silica minerals. This occurs
in
areas
with alternating
wet and dry
periods.
Groundwater silcretes
are
massive, with sharp upper
and
lower boundaries,
and
lack pedogenic features such
as
nodules
and columns.
The
structures
and
fabrics
of
the parent material
are retained
in the
silcrete. They probably represent, silicifica-
tion
on or
close
to the
water table
and are
related
to the
position
of the
paleowater table.
The
silica
is
imported either
by infiltration through
the
upper part
of the
profile
or by
lateral groundwater flow.
The
mixing zone between meteoric
ground water
and
saline waters from
a
lake
or the sea
facilitates silification processes
and the
silica rocks formed
there
are
considered
to be
silcretes
by
some authors.
Characteristics
of the two
types
of
silcretes, vaclose
and
groundwater, sometimes appear together because silicification
might have commenced with groundwater
and
continued into
the unsaturated zone (Bustillo
and
Bustillo, 2000). Generally
the silica precipitation
is
dependent upon geochemical para-
meters
at the
microscale
and the
pore water movement
at the
mesoscale.
The
movement
of
water
is
largely conditioned
by
geomorphological
and
hydrological controls,
and to
climati-
cally influenced water-table fluctuations.
Siicretes and climate
Varied studies
of
silcretes
in
Australia (Langford-Smith,
1978)
revealed these occur
in two
geological contexts:
(I)
associated
with deep-weathering profiles where their formation
is
attributed
to
wetter conditions;
and (2)
arid/semiarid
or
alkaline conditions.
In the
first context silcretes
can
coexist
with ferricretes
and
laterites
and in the
second with calcretes,
dolocretes
and
gypcretes. Within deep weathering profiles,
chemical alteration breaks
the
silicate minerals down
and
releases
the
silica.
The
mechanism
of
silica precipitation
is unclear
but
sufficiently sluggish drainage could lead
to
silica
precipitation.
In
arid/semiarid conditions,
the
source
of
silica
is
sometimes
a
consequence
of
the high amount
of
silica liberated
when other duricrusts (calcretes, dolocretes,
and
gypcretes)
are
formed
on a
silicate parent material. The dissolution
of
opaline
particles (phytolites, diatoms...)
and
eolic quartz sands
and
silica-rich groundwaters
are
other sources. Silica precipitation
mainly occurs
in
response
to
concentration through evapora-
tion
or by a
localized decrease
in pH.
The existence
of two
distinct climatic formation environ-
ments, does
not
mean
all
silcretes should
be
assigned
to
them.
Moreover, efficient geochemical
and
petrological tools capable
of definitively assigning
a
silcrete
to one of
these
two
contexts
have
not yet
been developed. After many years
use,
criteria
such
as the
amount
of
TiO2
or
some petrographic character-
istics,
are not now
considered acceptable (Nash etal., 1994).
Oxygen isotopic studies
of the
quartz that forms silcretes
may yield environmental information, including variations
in
temperature (Knauth, 1992).
The
opaline phases
and to a
lesser extent,
the
chalcedonic quartz,
are not
appropriate
for
the isotopic studies
due
mainly
to
bound/included water. Once
we know
the
oxygen isotopic composition
of the
quartz,
the
oxygen isotopic composition
of
fluids from which
it
precipi-
tated
can be
calculated
if the
temperature range
can be
established.
For
this calculation,
a
fractionation formula
can
be used. Webb
and
Golding (1998) consider
the
Knauth
and
Epstein formula gives more realistic results under near surface
conditions, which
is
where silcretes
are
formed. Detailed
geochemical
and
oxygen isotopic studies
of a
silcrete
and its
parent material
can
shed light upon
the
mechanism
of
silcrete
formation
and if
evaporation
is a
major factor
in
silcrete
formation (Webb
and
Golding, 1998). Silcrete formation
depends
on a
complex interaction between climate
and
silica
supply
and
needs tectonic stability.
M" Angeles Bustillo
Bibliography
Bustillo, M''.A.,
and
Bustillo,
M.,
2000. Miocene silcretes
in
argillaceous playa deposits, Madrid Basin, Spain; petrological
and geochemical features. Sedimentology, 47: 1023-1039.
Heaney,
P.,
1995. Moganite
as an
indicator
for
vanished evapoHtes:
a
testament reborn? Journalof Sedimentary Research, A65: 633-638.
Hesse,
R.,
1990. Silica diagenesis: origin
of
inorganic
and
replacement
cherts.
In
Mcllreath,
A., and
Morrow,
D.W.
(eds.), Diagenesis.
,
Canada: Geoscience Reprint Series, Volume
4, pp.
253-275.
Knauth,
L.P., 1992.
Origin
and
diagenesis
of
cherts:
an
isotopic
perspective.
In
Clauer,
N., and
Chaudhuri,
S.
(eds.), Isolopic
Signatures
and
Sedimentary Records. Berlin: Springer-Verlag,
pp.
123-152.
Langford-Smith,
T.,
1978. Silcrete
in
Australia. Australia: Department
of Geography, University
of
New England.
Nash,
D.J.,
Thomas, D.SG.,
and
Shaw,
P.A., 1994.
Siliceous
duricrusts
as
palaeoclimatic indicators: evidence from
the
Kalahari
Desert
of
Botswana. Palaeogeography. Palaeoclimatology, Palaeo-
eeology, 112: 279-295.
Thiry,
M., and
Milnes,
A.R., 1991.
Pedogenic
and
groundwater
silcretes
at
Stuart Creek opal field, south Australia. Journal
of
Sedimentary
Petrology,
61: 111-127.
Summerfield, M.A., 1983. Silcrete.
In
Goudie, A.S.,
and
Pye,
K.
(eds.),
Cliemieal Sediments and Geomorplwlogv. London: Academic Press,
pp.
59-61.
Webb,
J.A., and
Golding,
S.D.,
1998. Geochemical mass-balance
and
oxygen-isotope constraints
on
silcrete formation
and its
paleocli-
matic implications
in
southern Australia. Journal
of
Sedimentary
Re.search,
68A: 981-993.
Cross-references
Laterites
Siliceous Sediments
SILICEOUS SEDIMENTS
Introduction
Siliceous sediments
are
composed
of
silica that
has
actually
precipitated
at or
near
the
site
of
deposition
or has
replaced
pre-existing sediments. They
are
distinguished from clastic
or
terrigeneous sediments which
are
made
of
grains derived from
rocks elsewhere
and
physically transported
to the
site
of
deposition.
In
today's oceans,
the
dominant examples
of
siliceous sediments
are
oozes composed
of
microscopic silica
particles precipitated biologically
by
diatoms
and, to a
lesser
.SU.ICHJUS
SEDIMENTS
extent,
by radiolarians.
sponges,
and silicoflagellaies. Diatom
silica
is precipitated on a colossal scale in surface waters and
slowly
scltles to the ocean floor to prodtice diatomaceous
oo/cs.
Although most of ihe opaline tests dissolve in transit, a
small
peiceiiiage sur\ive and accunnilate in thicknesses Lip ui
hundreds
of
meters.
The purity of the ooze is greatest for
depths
below the carbonate compensation depth and in areas
far
removed from landmasses that can supply abundant
clays.
Siliceous
oozes are regularly encountered on the deep ocean
floor
and in deep-sea drill
cores.
Locally, radiolarians and
silica
sponges are so abundant that oozes of radiolarian
skeletons
and opaline spicules can also accumulate.
Diatomites, porcelanites, and bedded cherts
Along the coast of California, the Miocene Monterey
Formation has been uplifted to produce superb exposures of
diatomaceous layers displaying every gradation from pure.
unconsolidated diatomite to porcelain-like beds and even
brittle, glass-like beds of chert. The diatomite is extensively
mined for insulaiion. paint filler, filtration material, and a wide
\aricly of other industrial uses (Figure S23). The classic study
by Bramlette (1946) first demonstrated that the gradations in
lithology represented \arious steps in the burial transforma-
tion of opaline diatoms into microcrystalline quartz. Uplifted
exposures along coastal California allowed detailed examina-
tion of this process where it had been arrested in various states
of progress. Subsequent mineralogic and isotopic studies
revealed that the solution-resistant particles that had accumu-
lated continued lo resist dissolution during burial until
elevaled temperatures provoked an
itt.sitii
sohition-reprecipita-
tion conversion into hard, porcelain-like beds (Figure S24) of
hydrous opal with weak X-ray diffraction peaks of cristobafiie
and trydimite (opal-CT). Upon further burial, this opal-CT
converted to j:-quartz. Pore spaces and fractures that de\el-
oped at various times during burial became filled directly with
microcrystailine quartz or with fibrous quart-^ that grades
outward into microcrystalline aggregates of single, well-
formed crystals. The most recent stLidies (Beh! and Garrison.
19^4) suggest that much of the quartz in the Monterey
F'ormation formed throughout the burial history rather than
only at maximum burial as oiiginally thought. At burial
depths such that temperatures reach 80 100 C. all opal and
opal-CT becomes transformed into microcrystalline quartz.
Most ofthe great oil fields in California have developed where
abundant organic matter in diatomite converted to oil and
accLinuilated in fVactured masses of brittle porcelanite and
chert.
The stcpwise transformation so clearly displayed in the
Monterey Formation apparently also operates during burial of
all deep sea oo/es and readily explains the lithologic/
mineralogic variations of siliceous oozes encountered during
deep-sea drilling. Diatoms have become particularly abundant
in the last 30 million years and now dominate silica deposition
in the oceans. Tertiary examples of Monterey-type diatomitcs.
porcelanites. and eherts are particulary common in the circum-
Pacific area and have received much attention. Examples of
Eocene radiolarian oo/es that transf\)rmcd into porcefanite
and bedded ehert are also regularly encountered dLiring deep
drilling ofthe ocean floor.
Early observers noted peculiar regional beds of Paleozoic
and Mesozoic strata composed of interlocking grains of
microquartz that were clearly not clastic deposits. Examples
include bedded cherts in the Alps (Gruneau. 1965) and the
Devonian Arkansas/Caballos novaculite in Texas/Arkansas
(Miser and Purdue. 1929: I'olk and McBride. 1976). In many
fold belts, the cherts display a distinctive "pinch and swell"
texture and are known as "ribbon cherts". Chemical etching of
the microquartz now reveals radiolarian forms confirming
prior inferences that most of these bedded cherts originated as
radiolarian opal. The etched-out forms are distinct enough for
assigning biostratigraphic ages useful in tectonic studies.
Apparently, at times and in \arious places, radiolarians were
abundant enough to accLimulate as opaline oozes similar to the
Tertiary diaEomaceous examples. It is now inferred that these
transformed into bedded cherts throLigh the stepwise opal to
opal-CT to quartz sequence so beautifully documented for the
Miocene Monterey Formation.
Novaculite is a white radiolarian chert made almost
exelusively of intergrown grains of microquartz and is famous
for its widespread use as a whetstone. Other variations in color
and texture of radiolarian cherts probably relate to the clay
mineral content. The natLire of the depositional basins
combined with dilution effects of shale and calcareous
plankton and possible diagenetic segregation of silica from
shale (Murray etal.. 1992) may all play a role in determining
Figure S23 Qu<irry in Monterey Dialomite. Lompoc, Californi.i.
Figure S24 rtaiisitiun ol Lmiindlfd oigiiniL-iiLli Morileiey tliiitoniilf li
()|).il-(.T. I iammer head is on opal-CT hori^uns. Beatli cfifl at Pt.
Pedernales, California.
662
SILICEOUS SEDIMENTS
the Ihuil lithologic character of ihc uhcri. Inasmuch as ihc
mobile seafloor carries all sediments into accretionary wedges
and subduction zones within about 100 million years, most of
these older bedded cherts arc today found in highly deformed
and altered mountain belts.
Silicification
Silica is soluble in aqueous solution as H4SiO4. At pH>9. these
units polymerize and solubility of all solid forms of silica
increases dramatically. At lower pH and room temperature,
solubility of quariz is about 6ppm and opal is about 120ppm.
The solubility increases greatly with temperature, but because
of the long times avaiiabie. the low-temperature values are
high enough to allow large-scale transfer of silica in the pore
fluids of sediments prior to compaction. Supersaturation with
respect lo quarlz is achieved if ground waler moves through
uiicompacted sedimeni containing opal or grains of more
soluble silica such as volcanic glass and/or clay minerals.
Siiicitication can then occur if some other phase such as
metastable aragonite/calcite shell fragment (or a tree trunk
buried in volcanic ash) slowly dissolves or disintegrates and is
transported away by the through-going, quartz-saturated
ground water. Quartz can grow in the microscopic pore spaces
so vacated and replaeement can occur on such an incremental
basis that e\traordinary preservation of the precursor material
is achieved. Even forms as small as bacteria can be preserved in
this manner. Fibrous quart/ and small quartz crystals can also
grow in the larger void spaces from such saturated solulions.
Quartz crystals up to lOjim have been grown in seawater at
room temperature in 2 years in solutions with low silica
concentration (Mackenzie and Gees. 1971). If dissolved siliea
concentrations are >120ppm. silicification by opal can occur.
The long time-scales required to aehieve such incremental silica
replacement preclude realistic labi>ratory simulations; the
siliciiication process must be inferred from studies of the rock
record.
Mineralogy
Silica occurs in sedimentary rocks as several varieties of opal
and quartz. Opal is the most common form in Cenozoic
deposits, but older strata are always composed of quartz.
Amorphous, hydrous opal (opal-A) is today precipitated by
diatoms, radiolarians, sponges, and silicoilagellates. and also
forms inorganically in unusual depositionai environments and
in pore spaces during shallow burial. This material transforms
to quartz with increased time, temperature, and/or depth of
burial. Both nodular and bedded cherts are composed of
microcrystalline quartz with individual grains ranging in size
from about 8 20p (Figure S25) . The grains are thoroughly
intergrown with each other and this makes chert one of the
hardest of ihe sedimentary rocks. Each individual grain is
crystallographically complex, wiih many lattice faults and up
to
1
wt. percent OH and H^O. Fibrous quartz is typically
associated with the micro quartz (Figure S25) and has an even
more eomplex internal structure. Coarser erystals of quartz
with well-defined grain boundaries and more normal internal
crystallographic ordering often occur together with fibrous and
microcrystalline quartz and also form as recrystallization or
annealing products of niicroquartz.
Figure S25 Thin-scclion of chert under crossed nicols. Microquartz on
right hiilf, 2 bundles of fibrous qu<5rlz grade lo letl to coarser individual
crystals of quartz. From 2.7Ca silitified stromatolite in Forlesque
Group, W. Australia. Scale bar is l()|.im.
Nodular replacement cherts
The earliest examples of silica in sediments to be studied were
the nodular cherts in limestones (Figure S26). Famous as the
source rock for anthropological artifacts and widely used as
Hints for starling fires and igniting munitions (even used as
cannonballs in one battle of the American Civil War), nodular
cherts were originally interpreted as balls of opaline gel that
rolled around on the seafloor and hardened into quartz during
burial (Tarr. 1917). A long-standing debate started wheti Van
Tuyl (1918) argued that the nodules were actually replacement
features where zones of limestone had been incrementally
replaced by silica precipitated from later diagenetic fluids
moving through the sediment. Among others. Biggs (1957)
provided thorough documentation for early replacement and
the basic idea has not been challenged since.
Replacement of carbonate to produce nodular chcil occurs
during early sediment burial when shell friigments and other
primary grains of carbonate undergo dissoltition and re-
precipiiation into interlocking calcite and/or dolomite crystals
to Ibrm limestone and/or dolostone (Knauth, 1979. 1994). In
Phanerozoic rocks, the most common silica source is opaline
sponge spicules which dissolve and re-prccipitate as micro-
quart/ dtn"ing this transformation of sediment into sedimen-
tary rock. The amount of chert that forms depends on the
abundance oi' the silica secreting sponges in the original
depositionai environment and the vagaries of hydrologic
movement of pore fluids during stabilization. The nodules
are often among the first stable phases to form, and cross-
bedding, oolitic texture, and even organic-walled microfossils
are often exquisitely preserved. Primary textures in the
surrounding carbonate are typically degraded or obliterated
during this early diagenetic "stabilization" event making
cherts important ""windows" so the original dcpositonal
SILK FOLiS SEDIMENTS
Figure S26 Nodular rcpl.Kement chert in 1.2 Ga Bass Limestone,
Grand Canyon, Arizona.
siruclures and fabrics that have been otherwise lost
(Figure S27). This aspect of cherts has been generally under-
iitih/cd and offers much potential In the invcsiigiUiori of
carbonate depositional environments.
In addition to the proininent nodules, replacement ehcrt
iilso occtirs as small blebs, thin siringcis. silicilicd fossils, iind
disseminated patches throughoiit the parent carbonate mass.
The color of the chert is black if unieplaced impurities ;irc
organic matter, gray if remnant carbonate, and red if iron
oxide. The purer cherts Eire translucent and display eonchoidal
fracture (Kigure S2X). If pieces of such eheri are rciissembied,
they often do not fit back together perfectly. Apparently, there
is a slight ticxinij of ihe sample after breakage. Such pieces may
also luminesce when rubbed together m the dark. Other cherts
ha\e splinterey fracture and this usually indicates low-grade
metamorphic recrystalli/ation. Metamorphism can transform
the purer cherts into interlocking grains of coarse quartz that
are among the hardest known rocks. In thin section, cherts
display a large range of petrographic textures that relate to the
manner in whieh they formed and their subsequent burial and
iiietamorphic history (Hesse. 1990; Caro//i. 1993)
While samples of chert from the geologic column may
appear somewhat similar with regard to lithology and
petrographic texture, the oxygen isotope variation in them is
larger ihan (or any other group of terrestrial rocks (Knauth.
1992).
The "*O/'"O ratio is determined by the temperature of
Figure S27 Noduiar cherts displaying primary pisolitic iabrk which
was largely obliterated in hosi dolnslone by later neomorphism. Chuar
Formation,
Grand Canyon, Arizona.
Figure S28 t^ark nodular chert displaying Iranslucente and
conthoidal tracturo typical ot" near ll)()7,i replacement of limestone
with microquartz. Cretaceous Edwards Formation, Texas.
enccnintered in these rocks. One implication is that nodular
cherts tbrmed most abundantly in coastal en\i!-onments where
the transforniatioii of primary metastable grains to stable
crystals occurred in meteoric ground waters mixed with marine
pore fluids. Archean examples are all enormously depleted
quartz precipitation and by the 0/ O ratio of the pore in O relative to Phanerozoic examples, and it has been
fluid. Because there are 2 variables, there has been much argued that this indicates that the temperature of stabilization
eontrovers_\ over the origin of the enormous isotopic variation of initially deposited precipitates (and thus the climaiic
664
SILICEOUS SEDIMENTS
temperature) was up to 40 C warmer in the Archean (Knanth
and Lowe, 2003).
Replacement cherts are found in numerous Precambrian
carbonates where they can display well-preserved microtbssils.
Pine laminations in many Precambrian limestones and
dolostones are commonly defined by thin siliceous laminae
composed of microquartz. clay, and detrital quartz. These are
often classified as "stromatolitic"" layers formed by trapping on
biologic mats or precipitation of metastable carbonate grains
which are subsequently stabilized as limestone, dolostone, or
chert. The chert can form classic nodules, but more typically is
arrayed as thin discontinuous bands, blebs, and coarse laminae
(Figure S26). Black, dull-luster cherts in Precambrian stroma-
tolites have provided tnuch, if not most, of the Precambrian
organic-walled tnicrofossils that record the paleontologic
record of early life. The silica source lor these eherts retiiains
problematic. Ocean water itself almost certainly had much
higher dissolved silica in the absence of abundant silica
secreting ot"ganisms to remove it and also due, possibly, to
higher saturation values from much wanner climatic tempera-
tures.
It is also possible that some varieties of Precambrian
bacteria precipitated silica (atid carbonate) directly. In any
case,
prc-compaetion silica diagenesis was extremely active.
and distinct zones and laminae were often replaced with
microquartz before bacteria could degrade or decompose.
Although nodular replacement cherts in carbonates are
comtnon in rocks older than mid-Tcrliary, they have never
been found forming anywhere today. The "disappearance'" of
the process of nodular chert formation is roughly coincident
with the explosive radiation of diatoms in the mid-Tertiary and
with the apparent reduction of silica sponge populations in
shallow waters. Silica precipitation by diatoms is so prolilic
today that dissolved silica in surface and shallow waters is
usually <5 ppm. Opaline sponge spictilc debris is still produced
in shallow waters in relatively minor amounts, but the spicules
appear to quickly re-dissolve in the undersaturated. shallow
seawater and do not contribute significant dissolved silica for
diagcneiie pore fluids. It seems that the advent ol'diatoms so
reduced dissolved silica concetitrations in shallow waters that
opaline sponge spicules can no longer accutnulate in shallow
water carbonate depositional environtnents. The inability to
observe eherts forming today means that all knowledge of
replacement cherts tnust presently be based oti inferences
deduced from investigations ofthe rock record.
In addition to the nodular replacetnent cherts and the
bedded cherts, there are nimierous other exatnples of silica
accumulations and replacetnents that occur during early
diagenesis in sediments of all types. These included silicified
fossils of everything from bacteria to trees, interbedded layers
of chert and iron oxide (iron formations), silicified evaporites,
silicified voicaniciastic materials, and vein/cavity in-filis in all
types of seditnentary and volcanic rocks. The processes and
timing of silica diagenesis in many of these exatnples remain
poorly understood.
Silicified evaporites
Soluble evaporite minerals such as gypsutn. anhydrite, and
halite often oecur in the supratidal facies of litnestoncs and
dolostones. These phases are the first to dissolve during early
diagenesis when coastal ground waters containitig a tneteoric
water component percolate through the original marine
precipitates. Vugs and tnolds are thus encountered by
silica-saturated percolating groundwaters and can becotne
filled with quart/, flbroits quartz, and/or microquart/ . Small.
ctii-sized cauliflower-shaped nodules of chert that tnitnic the
original morphology of gypsum/anhydrite tnasscs arc thus
cotninoii in supratidal carbonates (Chowns and Elkins. 1974).
The surrounding carbonate groundmass is subsequently
dolomitized. but the distinctively shaped chert nodules provide
instantaneous recognition of a probable supratidal deposi-
tional environment. Even in cases where chert nodules also
replace the areas immediately surrounding evaporite crystals,
coarse quartz in-fills of dove-tailed twins of gypsum or even
halite cubes can be observed in petrographic thin-section.
These examples of silicification provide powerful indicators of
the supratidal environment in carbotiate roeks so old that all
traces of the pritnary evaporitic minerals have been retnoved
by later dissolution.
Magadi cherts
The youngest cherts are probably those that fortned in and
around Pleistocene lakes in the rift valleys of East Africa.
There, hydrous Na-silicates such as magadiite arc precipitated
in varve-like layers in alkaline lakes whieh undergo seasonal
evaporation. Grid-work arrays of submicroscopic quartz
crystals apparently replace the Na-silicate phases prior to
cotnpaction (Schubel and Sitnonson. 1990). Oxygen isotope
studies suggest that the diagenetic fluids had become cvapora-
tively enriehed in dissolved silica. Silicilication of trona and
other lacustrine cvaporatite tnitierals are also evidetit. Several
Tertiary examples of cherts that fortned in this manner have
been identified (Sheppard and Gude. 1986).
Precambrian cherts
Bedded cherts and silicified strata of all kinds are widespread
in Archean Greenstone Belts. Silica was apparently vastly
tnore tnobile in these depositional environtnents and early
silicification of sediments appears to have occurred on a scale
not observed since. A possible explanation is that dissolved
silica was fed into the deep oceans by hydrothermal vents on a
vast scale and that silicification in the vent areas was intense.
However, massive silicification does not appear to occur
around modern examples and modern deep sea hydrothermal
silica deposits are limited to local precipitation of opaline
particles ("'white smokers") induced by the sudden encounter
of silica-saturated hydrothertiial water with cold ocean water.
It is possible that the Archean climates were dratnatically
wartner atid that the Archean oceans therefore had dt"atnati-
cally higher silica solubility. Silicification in hydrothermal cells
is thereby facilitated as more silica-rich ocean water is
circulated through the adjacent sediment pile. More extensive
deposits of opaline floes and particles around hydrothertnal
vents tiiay have been widespread, and these would have
converted to quartz chert during early diagenesis.
Also abundatit in the Archean as well as in the early
Paleoproterozoic arc the famous and etiigmatic "'Banded iron
Formations" of interbedded chert and iron oxide. Peaking In
abundance at prior to 2.0 Ga. these also fortned on stable
platforms in water shallow enough for stromatolites produced
by photosynthetic cyanobacteria to occasionally accutnulate.
The iron in these deposits was apparently itijected as Fc^" into
deep anoxic ocean water from far-away hydrothertnal vents.
The dissolved Fe^" was oxidized in surface waters and/or in