Category Archives: Under the microscope

Miller indices in crystallography

Minerals are defined by their chemical composition and their crystal forms. Miller, and Miller-Bravais indices are the standard where every face on a crystal is given a unique description that in general notation is written as (hkl) and (hkil). Two of the most common forms are prisms (the tourmaline crystal on the left), and pyramids as shown in the tourmaline crystal termination and the volcanic quartz.

Miller, and Miller-Bravais indices are the standard where every face on a crystal is given a unique description.

Crystals are structured solids made up of an ordered arrangement of ions, atoms, or molecules. At the atomic scale this three-dimensional array is called a lattice. The smallest, and most basic lattice representation of any crystal is its unit cell. The shape and symmetry of the crystals we observe and the arrangement of their faces, mimic the geometric properties of their unit cells.

The type of structure we see, specifically the outward appearing form of the crystal (shape, size, symmetry), depends on several factors:

First and foremost, the composition of the ordered components; their atomic dimensions and charge (for cation and anions).
Environmental factors such as the concentration, activity, or fugacity of their constituent atoms, ions, or molecules in fluid, molten or gas phases,
Temperature and pressure conditions,
Biological mediation of precipitation at Earth’s surface, for example aragonite in seawater whitings, ooids, and microbialites.
Kinetic factors that determine the rate of chemical reactions and crystal formation (precipitation). A classic, and still problematic example in sedimentary geology is the precipitation of dolomite.

Crystallography was in its infancy in the late 18^th and early 19^th centuries. An important advancement at this time by Christian Weiss (1780 – 1856) was the recognition that crystals could be thought of as collections of similar planes arranged symmetrically around three axes. Weiss devised the three-axis system a, b, c, and determined that any single crystal face will intersect one, two or all three axes. Other faces of the same crystal could then be described according to the proportions of their intersections with the three axes.

William Hallowes Miller (1801-1880) took this a step further. Miller was a Welsh mineralogist who eventually assumed the role of Professor of Mineralogy at Cambridge University (replacing William Whewell). His contribution, published in A Treatise on Crystallography (1834), applied spherical trigonometry to a numerical system of crystal nomenclature; Millers systematic approach stands today.

Crystal axes, crystal systems, and crystal face intersections

Definition of the seven crystal systems is based on the relative lengths of the axes (labelled a, b, c) and their angular relationships. Five systems have three axes and two have four axes (hexagonal, trigonal):

Isometric; a = b = c, all at 90^o. This class has the greatest degree of symmetry.
Tetragonal: a₁ = a₂≠ c, all at 90^o.
Orthorhombic: a ≠ b ≠ c, all at 90^o.
Monoclinic: a ≠ b ≠ c, ab = bc = 90^o, ac ≠ 90^o.
Triclinic: a ≠ b ≠ c, ac ≠ bc ≠ ab ≠ 90^o.
Hexagonal: a1 = a2 = a3 at 120 ^o, ≠ c at 90^o. One 6-fold axis of symmetry.
Trigonal: a1 = a2 = a3 at 120 ^o, ≠ c at 90^o. One 3-fold axis of symmetry.

Note that we label each axis according to whether it is positive or negative; by convention, the negative labels place the minus sign above the letter.

Eighteenth and 19^th C crystallographers did not have the luxury of X-ray radiography to identify unit cells. Instead, they described crystal forms by measuring the interfacial angles of real crystals and determining their relative intersections with the crystal axes. The phrase ‘relative length’ is important; identification of a class does not depend on some standard axial length or crystal size. This also means that the intersection of a crystal face with any of the axes will also be relative; if a face moves parallel to itself, the relative intersections will remain the same. Every crystal face will intersect one, two or three axes.

Christian Weiss crystal face parameters

The crystal face labeling system Christian Weiss devised was based on the relative axis intersections determined by measuring interfacial angles. If we are to describe all faces comprising a crystal, we need a standard face against which all other faces are measured. The intersections on this standard face are assumed to have a value of one.

Intersection with one axis

For crystals having faces that intersect only one axis, the intersection value is assigned 1; intersections on the other two axes are at infinity. For example, if the face intersects the c axis it will be labelled ∞a, ∞b, 1c. (∞ means intersection of an axis at infinity).

Examples of crystal faces intersecting one and two axes. The isometric system crystal is a cube of equal sides and equal axis intersections (left). The prism is another common crystal form that in this case intersects two axes.

Intersection with two axes

Likewise, for crystals with faces intersecting two axes, one of the faces is assigned the value 1, and the other will be a fraction (or multiple) of one. The third axis intersection is infinity. In the example shown below, the prism face notation is 1a, 1b, ∞c.

Intersection with three axes

Faces having three intersections are treated in the same way, but in this case the largest face on any crystal that intersects all three axes (not necessarily the largest face on the crystal) is chosen as the unit face. The intersections are all given a value of unity, so that the unit face can be described as 1a, 1b, 1c. All other faces in the same crystal can now be described with reference to the unit face.

The orthorhombic crystal shown here consists of a prism and two sets of pyramid terminations (each set is a mirror image). All faces in the pyramids intersect the three axes, but the faces on the middle pyramid are the largest so we assign one of these as the unit face. The values of the three intersections on the unit face are 1, so the face is labelled 1a, 1b, 1c.

With the unit face as our standard for this crystal, we can now tackle the other faces. Note that the actual intersection point may need to be extrapolated. For the small set of pyramid faces the intersections with ‘a’ and ‘b’ are extrapolated – their (approximate) values relative to the unit face are indicated. Thus, the face can be labelled 2a, 2b, 2/3c. Because these values are relative, we can divide by 2, so that the terminal pyramid face can be assigned 1a, 1b, 1/3c.

A schematic of crystal face intersections with two and three axes in an orthorhombic crystal, from which Weiss intersection ratios and Miller indices are calculated. The unit face is the large pyramid face that intersects all three axes (diagram centre).

William Miller intercedes

Weiss’ method is certainly workable, but it is regarded as cumbersome. Miller’s modification was quite simple, a simplicity that has survived. It has four parts:

Identify the Weiss parameters for each face.
Determine the reciprocal for each axis value.
Normalize the values so that each face is described by integers only (i.e., clear the fractions).
The axes are always described in the order a, b, and c; therefore, the label can omit these. Convention also requires the Miller Indices to be enclosed in brackets.

For crystal systems with three axes, the general notation for any crystal face is (hkl).

For the examples shown above, the reciprocal of the single face intersection with the axis c becomes:

1/∞a, 1/∞b, 1/1c which simplifies to 0a, 0b, 1c, or in the Miller indices convention, (001).

For the two axis intersections (tetragonal prism), the reciprocal is 1/1a, 1/1b, 1/∞c and the Miller designation (110).

For the orthorhombic crystal, the reciprocals of the unit face intersections are 1/1a, 1/1b, 1/1c, and the Miller designation is (111); for the smaller pyramid face (113), and the prism face (110). For axis intersections that are negative, the number has a minus sign above the integer.

Miller indices for crystals having faces that intersect one (left), two, and three axes (right).

Miller-Bravais indices

The hexagonal/trigonal groups of crystals are the odd ones out in the crystal classification system; they have four axes, three equal axes at 120^o labelled a1, a2, and a3, that are at right angles to the c axis. Trigonal crystals are frequently described in terms of hexagonal geometric properties. The general form of a hexagonal crystal face is (hkil). Quartz is the ubiquitous mineral in this group, commonly forming prisms with pyramid terminations and in some cases, bipyramids like the example shown below (this quartz habit is common in felsic volcanic rocks). Labeling hexagonal crystal faces follows the same method for both the Weiss and Miller indices.

There is an additional rule that the sum of h + k + i = 0. Prism faces that intersect two of the ‘a’ axes but not the c axis have Weiss parameters like 1a_1, ∞a₂, -1a₃, ∞a₄, and a Miller-Bravais index like (101^–0). Likewise, pyramid faces that intersect two ‘a’ axes and the ‘c’ axis will have indices like (101^–1) (the sum of hki is 0). For crystals having pyramid faces that intersect 3 ‘a’ axes, the middle axis intersection will be a fraction of that for the other two.

Miller-Bravais notation for pyramid faces intersecting two ‘a’ axes and the c axis in a slightly abraded grain of bipyramidal quartz (scanning electron micrograph).

Twin and cleavage planes

The Miller and Miller-Bravais notations are also applicable to the description of cleavage and twin planes because they generally parallel certain crystal faces. For example, the prominent rhombohedral cleavage in calcite is (1011), in gypsum the perfect cleavage is (010) and less perfect (100) and (111). The prominent basal cleavage in biotite and muscovite is (001) – the common form of detrital mica flakes in sediments. Albite twinning is across the (010) plane – this can also be a cleavage plane, along with (001).

A couple of good resources

Online mineralogy lectures by Prof. Stephen A. Nelson at Tulane University

Introduction to mineralogy: Tark Hamilton, Camosun College.

There are lots of YouTube presentations on the topic.

A geological life

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Negotiating a passage through remnants of sea ice, beginning the 1977 field season on Belcher Islands.

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Not really a CV, more a precis

1969 and about to begin a BSc at Auckland University, Aotearoa-New Zealand (aiming to major in chemistry), I needed one more paper to complete the first-year syllabus. A friend suggested I try geology – “isn’t that fossils, rocks, dirt and stuff?” “Yeah, pretty interesting though”. “That’ll do!” I had made my choice. At that time BSc geology consisted of 2 papers each year – one paper (imaginatively called Geology A or B – I can’t remember which) included all ‘soft rock’ topics (sedimentology, paleontology, stratigraphy, geophysics etc.), and the other paper mostly ‘hard rock’ topics (igneous, metamorphic, structure, crystallography etc.). Students were immersed in everything – there were no options.

So, I majored in geology and chemistry for my BSc (Auckland University), but my passion was geology.

Fast forward another 2 years and completion of an MSc (geology, Auckland University, 1975), with a thesis on the sedimentology and stratigraphy of Quaternary aeolian, shallow marine, and fluvial deposits in subtropical northernmost New Zealand. Much of the exposure was coastal so in between measured sections and samples, I collected shellfish or threw out a fishing line to catch dinner.

From 2 million years to 2 billion years: It’s now January 1976, confronted by a minus 25^oC Ottawa, Canada, on my way to start a PhD at Carleton University. I’d never seen so much snow. My supervisor was to be Alan Donaldson, a Precambrian seds guru, who was pretty keen on me doing a thesis on the sedimentology and stratigraphy of a Paleoproterozoic succession on Belcher Islands, Hudson Bay. This collection of elongate, squiggly islands is held up by a 7-9 km thick, 1.8 to 2 billion year-old succession of stromatolitic platform carbonates, shallow marine siliciclastics, a banded iron formation, a turbidite succession, red beds, and two spectacular volcanic successions. When I asked Al which formation I should work on he said “do the whole shebang“. Ok! A tough working environment but great exposure of fabulous rocks (lots of boat – Zodiac work over 0^oC seawater and inclement weather – my assistant and I wore life jackets so that, we were told, they could find our bodies for insurance purposes). Al was a great supervisor.

So over two 10-week field seasons (1976-77) I measured 20,000 m of section, 2000 paleocurrent directions, and a slew of petrographic analyses. I defended in June 1979.

Next stop Calgary for a two-year stint with Gulf Canada Resources, then in 1981 landed a job with the Geological Survey of Canada. In the GSC Calgary office most of that work was centred on field mapping and stratigraphy-sedimentology of Upper Cretaceous – Paleogene clastic sediments on Ellesmere and Axel Heiberg islands – mostly shallow marine, fluvial and delta settings. There were other bits and pieces in the Alberta Front Ranges, Yukon, and Mackenzie River. A move to the GSC Vancouver office in 1989 saw the emphasis change to Mesozoic sedimentology of one of the largest Intermontane basins in British Columbia, Bowser Basin, a foredeep with sediment derived from an obducted slice of oceanic crust.

Between 1992 and 1993 I realized I needed a change in Earth Science emphasis, and a logical choice was hydrogeology where I could use my knowledge of sedimentary rocks (sediment body geometry, composition, porosity-permeability and so on) and geological mapping to characterize fluid flow at both a basin-scale, and near-surface groundwater aquifer scale.

This culminated in a 4-year GSC pilot project to map and characterize aquifers in the greater Vancouver – Fraser Valley – Delta region of southern British Columbia, coordinating the expertise of colleagues with the use of shallow reflection seismic, ground penetrating radar, electromagnetics, gravity, MODFLOW modelling, and GIS data management of water-well databases and subsurface aquifer maps. Inserted between these programs was a brief secondment to the Hungarian Geological Survey in 1992 to help develop their basin analysis projects.

I quit the GSC in 1997 and moved with my Canadian family back to Aotearoa – New Zealand to work a 4-year teaching stint at University of Waikato, working primarily on Late Miocene – Pliocene siliciclastics and cool-water carbonates in Whanganui Basin (west North Island). From there a part-time position at Auckland University, teaching post-graduate basin analysis and undergraduate hydrogeology, and supervising (mostly) groundwater-related theses. But this position also required a significant weekly commute, and with an evolving disenchantment of academia I decided to form my own consulting company in 2005 – a company of one – and never looked back!

As a consultant I worked lots of NZ coal and oil-gas well-site geology (yes, I know…!), geothermal hydrogeology (Taupo region), basin-scale CO₂ sequestration evaluation (with GNS), lithium-bearing brines in the Chilean Altiplano (base camp at 4000 m – the geology here is amazing), and a bunch of smaller jobs mostly groundwater-related. I loved this stage of my life – I was my own boss!

Now I am retired, tending to our organic kiwifruit orchard – we’ve been at it 25 years (all fruit exported so keep an eye open for it on your supermarket shelves), and maintaining the geoscience website that you must have linked to because you are reading this.

So, almost 50 years as a geologist. Time during 40 of those years was also spent as Editor and Associate Editor, mostly for the Canadian Society of Petroleum Geology and the SEPM including 8 years on the SEPM Council, and reviewer for dozens of journal papers. I still get requests for paper reviews, but I politely decline, noting that it is someone else’s turn.

Would I do it all again? Damned right, although I would probably try to insert a chapter on planetary geology into that life.

Lithic grains in thin section – avoiding ambiguity

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Green schist hypothetically ‘eroded’ and whittled progressively to finer grain sizes

The utility of lithic grains as provenance indicators

The composition of terrigenous sandstone is usually described in terms of the tripartite quartz-feldspar-lithic framework. The quartz and feldspar clast end-members are monomineralic. The lithic end-member takes care of all other grain types that are a mix of minerals such as quartz, feldspar, clay, heavy minerals, volcanic groundmass, and carbonate.

Notwithstanding the value of zircon systematics to unravel potential sources of clastic sediment, most sandstone provenance studies begin with an evaluation of the granular framework. Unfortunately, the value of the common rock-forming minerals quartz and feldspar as provenance indicators is frequently ambiguous – a grain of monocrystalline quartz may be derived from granitic, rhyolitic, or gneissic source rocks; and polycrystalline quartz from several kinds of metamorphic rock, recrystallized sedimentary chert, or tectonized quartz-rich rocks. In comparison, a decent lithic is worth its weight in gold.

However, there are a couple of limitations that influence the value of lithic fragments as source rock indicators:

The potential for diagenetic alteration of some components, particularly clay, feldspar, ferromagnesian minerals, and volcanic glass, and
Grain size will influence the identifiability of source rock composition, texture, fabric, and structural attributes. As a general rule, larger clast sizes will afford greater confidence in interpretation.

The exercise below shows how the identification and interpretation of lithic clasts can change according to:

The crystallinity or coarseness of the original source rock.
The final grain size of lithic fragments.

Green schist

The first example is a quartz-rich green schist (Fox Glacier, New Zealand):

Foliation is defined by biotite laths and elongation of quartz crystals.
Mica crystal orientations are parallel from one foliation to the next.
Quartz crystals have irregular, interlocking contacts.
Quartz crystal size ranges from <100 μm to 600 μm (0.6 mm), corresponding to very fine through coarse sand size.

Wide-field view of a green schist (3.8 mm wide (3800 μm). The solid yellow outline corresponds to a hypothetical, very coarse sand-size clast. Dashed outlines correspond to medium sand-sized clasts. Crossed polars.

In the next image, hypothetical sand grains are ‘eroded’ from the green schist, beginning with the very coarse sand-sized fragment that, in turn, is whittled progressively to finer grain sizes (grain sizes are based on the Wentworth Scale).

Lithic fragments composed of green schist (a mix of quartz and biotite) at progressively finer grain sizes. Views are all crossed polars.

Very coarse to coarse sand; The full complement of interlocking quartz grains, biotite laths and schistose foliation is recognizable in both grain sizes.
Medium sand; quartz aggregates would probably be counted as polycrystalline quartz rather than lithic grains. Biotite may persist in medium sand-sized grains but its orientation at this scale of observation may not reflect the larger-scale foliation. Furthermore, the identification of foliation in quartz is also ambiguous.
Fine to very fine sand; In this example, fine and very fine grain sizes have the same dimensions as many individual quartz crystals. In this case, quartz grains might be counted as monocrystalline or polycrystalline quartz rather than as lithics.

Biotite-hornblende gneiss

Biotite-hornblende gneiss. The solid yellow outline corresponds to very coarse sand-size; the dashed outlines correspond to medium sand. The straight dashed line parallels the foliation. Crossed polars.

Gneissic rocks tend to be more coarsely crystalline than schist. In this example and at this scale (field of view is 3.8 mm wide), the foliation is recognizable but less distinct than in the green schist shown above. Biotite crystals are larger than their green schist counterparts and oriented at more variable angles to the foliation.

Lithic fragments composed of biotite-hornblende gneiss at progressively fine grain size. Views of the very coarse sand grain are crossed polars on the left, plain polarized light on the right. All other grains are crossed polars.

Very coarse sand; This grain size captures large, untwinned feldspar, quartz, plagioclase (albite twinning), and a couple of biotite laths. However, the likelihood of capturing the gneissic foliation is low, even in very coarse-grained clasts.
Coarse-grained and finer sand; Foliation is not identifiable in these grains.Medium sized grains may capture crystal a boundary, but many would appear monomineralic and difficult to differentiate from other sedimentary, volcanic, and intrusive source rocks.
If grains at very coarse sand or larger are not available, then the distinction between this gneiss and possible granite, granodiorite, or reworked sedimentary source rocks will be ambiguous.

Biotite granite

Biotite granite. The solid yellow outline corresponds to very coarse sand-size; the dashed outlines correspond to medium sand. Crossed polars.

The granite consists of untwinned potassium feldspar, quartz, plagioclase, and biotite. Crystal sizes range from 100 μm to more than 1500 μm (1.5 mm), averaging 400-600 μm. There is no foliation. There is little evidence of strain in the quartz crystals.

The following diagram shows the progressive change in grain size and identifiable features that would lead to an interpretation of granite source rock.

Lithic grains of granite composition at progressively fine grain size. Views of the very coarse sand grain are crossed polars on the left, plain polarized light on the right. All other grains are crossed polars.

Very coarse sand; Grains having this dimension have captured the essential feldspar-quartz-biotite aggregates. However, even at this grain size there may be difficulty distinguishing granite from some gneissic rocks, particularly if quartz shows any strained extinction.
Coarse sand; This grain size may capture single crystals of quartz, feldspar or ferromagnesian minerals, in which case they may not be distinguishable from gneiss, acid volcanic, or reworked coarse-grained arenites. This ambiguity will persist even if crystal aggregate boundaries are preserved.
Medium to very fine sand; Most grains will be counted as monomineralic. There will be little diagnostic information in these grains to distinguish them from many other source rock compositions.

Olivine basalt

Olivine basalt, where olivine (Ol) and plagioclase (Plag) phenocrysts occur within an aphanitic groundmass. The solid yellow outline corresponds to very coarse sand-size; the dashed outlines correspond to medium sand. Crossed polars.

Like many basalts and andesites, this example contains centimetre-scale phenocrysts surrounded by a sea of very small plagioclase laths in much finer-grained groundmass (glass or altered glass). Volcanic textures like this are very distinctive and the corresponding lithic fragments are perhaps the most easily recognizable of all lithic types.

Olivine basalt lithics at progressively fine grain size. Views of the very coarse sand grain are crossed polars on the left, plain polarized light on the right. All other grains are crossed polars.

Sand-sized clasts consisting of volcanic material are some of the more easily recognizable lithic grains, particularly if they contain groundmass.
Very coarse to coarse sand; Phenocrysts may be captured in the coarsest size fraction, but commonly will tend to separate from the groundmass during weathering and erosion.
Coarse sand and finer; Volcanic groundmass is characteristically glassy or altered glass that is anisotropic under crossed polars. Groundmass commonly contains a mass of very small plagioclase crystals; crystals may show some alignment (a product of original flow banding). These fabrics are common in basaltic and andesites rocks, and serve to distinguish them from mudrock lithic grains, even at fine sand size.
Diagenetic alteration of the groundmass may render some lithics indistinguishable from siltstone – mudstone sourced lithic grains.

Recrystallized chert

Part of a chert granule in an arkosic arenite. The long dimension of the grain is 1.5 mm. Crossed polars.

Part of a chert granule in an arkosic arenite. The arrow indicates radial fabrics developed from recrystallization of chert. The long dimension of the grain is 1.5 mm. Crossed polars.

This example looks at the breakdown of an existing chert lithic clast. There is a diffuse band of partly recrystallized silica where the crystal size changes from <10 μm to about 100 μm (0.1 mm). Recrystallization has also produced some radial silica clusters (arrow).

The micro-cryptocrystalline textures of typical chert lithic grains are relatively easy to distinguish from other quartz aggregates
They are recognizable at grain sizes at least to fine sand.
Grains consisting of the coarser recrystallized silica might pass for polycrystalline quartz grains unless radial crystal aggregates are also present.
Incomplete recrystallization commonly produces patchy or clotted textures where patches of remnant microcrystals grade into coarser crystal aggregates, reminiscent of the structure grumeleuse in neomorphosed micrites.

Silty mudstone

Silty mudstone with recognizable fine to coarse silt-sized quartz, chlorite, carbonate, and feldspar cleavage fragments. There is a hint of (horizontal) lamination. The matrix is a mix of clays, carbonate, and iron oxides. The field of view is 1.2 mm wide. Plain polarized light.

There is a complete gradation from very fine framework clasts to matrix. This is recognizable in grains at least to medium sand size.
Textures in some fine-grained sedimentary lithics may resemble the groundmass of aphanitic volcanic rocks, particularly if the feldspar laths have been diagenetically altered in the latter.
Sedimentary layering in mudrock lithics must be distinguish from flow alignment of feldspars in aphanitic volcanic rocks.

Some generalizations

The confidence with which clastic sediment source rocks can be identified depends strongly on the crystallinity or coarseness of potential sources. As a general rule, the coarser the crystallinity, the larger the clasts needed to positively identify that source.
The above comment can be restated as – there is an optimum grain size where the whole-rock identity of a lithic fragment can be made. For example, the optimum clast size that captures the essential minerals and foliation for schist-type rocks will be smaller than that for coarser gneissic rocks. At progressively finer grain sizes the common rock-forming minerals (quartz, feldspar, various heavy minerals) tend to be presented as monomineralic grains rather than components of lithic grains.
The specific source rock identity of common monomineralic, and even some polycrystalline aggregates of quartz and feldspar is frequently ambiguous, making the distinction among granite-granodiorite, gneiss, volcanic phenocrysts, or recrystallized chert difficulty. There are a few exceptions, such as bipyramidal quartz that is restricted to rhyolite, dacite, and ignimbrite, or minerals like kyanite and sillimanite that are found in high-grade metamorphic rocks.
For provenance studies, choose the coarsest samples possible.

Neomorphic textures in thin section

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Pervasive neomorphism of this oolite sand has increased calcite crystal size overall, and in the process has overprinted the primary concentric layering. Plain polarized light.

The term neomorphism is reserved for recrystallization of a mineral and its polymorphs to new crystal phases having the same composition. In calcium carbonates this involves recrystallization of calcite, high-magnesium calcite, and aragonite to calcite (usually low-magnesium calcite). Dolomites are also prone to neomorphism (dolomite to dolomite replacement); note that the replacement of calcite or aragonite by dolomite is not a neomorphic process. Neomorphism commonly results in an increase in crystal size and the destruction or overprinting of original fabrics.

The actual process of neomorphism is still a bit of mystery. Initiation of recrystallization begins with dissolution at the crystal-fluid interface. Subsequent precipitation must involve changes in the local thermodynamic conditions (such as increased saturation, or a change in the activity of carbonate species). If crystal size continues to increase (via syntaxial growth) then the transfer of dissolved mass to the crystal interface must continue.

Neomorphic versus cement textures

Cements grow by precipitating into the available inter- or intragranular pore space. Thus, crystal faces are free to grow and therefore they tend to be planar, with sharp crystal edges and terminations. In contrast, the locus of neomorphic replacement of pre-existing crystal masses (e.g., cements, bioclasts) and subsequent precipitation will tend to be focused along crystal discontinuities such as cleavage and twin planes, as well as crystal edges and terminations, in part because they are regions of higher surface free energy. Thus, neomorphic crystal boundaries are commonly irregular and characterised by embayments, re-entrants, or segmented.

Neomorphic replacement of carbonate cements, bioclasts, microbial laminates, and mudrocks commonly proceeds fastest in fine-grained deposits such as micrite – again this may be a function of higher surface free energies relative to much larger crystals and coarse-grained clasts.

Recrystallization may also show a progressive change in crystal size – referred to as aggrading neomorphism. Crystal aggradation can mimic pre-existing cement fabrics where crystal size increases from the cement boundary into pore spaces. Aggrading neomorphism is common in carbonate mudrocks; the resulting textures appear clotted, where remnant patches of micrite appear to float among clusters of coarser crystals. This kind of texture is also called structure grumeleuse after Lucien Cayeux. Clotted textures may also be pellet-like; the distinction between original peloidal textures and these neomorphic fabrics requires careful attention. Neomorphic fabrics frequently cut across cement and clast boundaries.

Some criteria that distinguish neomorphic from cement fabrics are listed below; the information is gleaned from R.L. Folk (1968), Robin G.C. Bathurst Chapter 12 (1976), J.W. Morse and F.T. Mackenzie 1990, and Erik Flugel. (2010, Chapter 7).

And in thin section…

Intercrystal boundaries in cements

Intragranular calcite cement in bryozoa pores https://www.geological-digressions.com/carbonates-in-thin-section-bryozoans/ (the much smaller, dark, circular bodies within the bryozoan structure are zooids). Cementation began with small, bladed calcite crystals lining pore walls, overlain by coarser, more-or-less equant spar. Most intercrystal boundaries are straight. Boundaries that appear curved actually consist of smaller, straight-edge segments. Left: Plain polarized light. Right: Crossed polars. Sample is from the Oligocene Te Kuiti Group.

Intragranular calcite cement in bryozoa pores (the much smaller, dark, circular bodies within the bryozoan structure are zooids). Cementation began with small, bladed calcite crystals lining pore walls, overlain by coarser, more-or-less equant spar. Most intercrystal boundaries are straight. Boundaries that appear curved actually consist of smaller, straight-edge segments. Left: Plain polarized light. Right: Crossed polars. Sample is from the Oligocene Te Kuiti Group.

Detailed view of cavity-filling cement fabrics in coarse calcite spar. Crystal boundaries are mostly straight (planar) and terminations are sharp. There is a gradation in crystal size from small, scalenohedral crystals lining the cavity wall (arrows, bottom of image) to much coarser spar in the cavity interior. Some of these smaller crystals may have developed incipient neomorphism. Left: Plain polarized light. Right: Crossed polars.

Neomorphic fabrics

A mix of cement and neomorphic textures. Inter- and intragranular calcite spar cements in Oligocene barnacle (ba) https://www.geological-digressions.com/carbonates-in-thin-section-echinoderms-and-barnacles/ – bryozoa (br) limestone. Some of the smaller crystals lining bioclast rims have scalenohedral terminations (back arrows). The bryozoan intragranular zooid pores are filled with dark brown micrite (white arrows), but in some the micrite has recrystallized to calcite spar (yellow arrows). In this case, crystal boundaries are less distinct because the neomorphic spar is intergrown with remnant micrite. Top: Plain polarized light. Bottom: Crossed polars.

A mix of cement and neomorphic textures.
Inter- and intragranular calcite spar cements in Oligocene barnacle (ba) – bryozoa (br) limestone. Some of the smaller crystals lining bioclast rims have scalenohedral terminations (back arrows). The bryozoan intragranular zooid pores are filled with dark brown micrite (white arrows), but in some the micrite has recrystallized to calcite spar (yellow arrows). In this case, crystal boundaries are less distinct because the neomorphic spar is intergrown with remnant micrite. Top: Plain polarized light. Bottom: Crossed polars.

Contrasting neomorphic and cement textures in a Jurassic bivalve.
Neomorphic fabrics in the (possibly aragonitic) shell preserve some of the original crossed-lamellar structures that parallel the shell wall (yellow arrow). Crystal boundaries are highly irregular, producing an irregular interlocking framework. In contrast, cementation of the shell interior (where the animal once resided) began with isopachous, scalenohedral calcite that was followed by coarse calcite spar. The skinny brown layer is the interior shell wall, now replaced by siderite. Crossed polars.

A mix of dolomite spar cement and finer neomorphic dolomite in a late Precambrian cryptalgal laminate, Tindir Group, Alaska. The large dolomite crystals preserve original crystal faces and terminations. The patches of finer dolomite crystals have boundaries that are more embayed or diffuse (arrows); the patches appear to float in the coarser fabrics. Left: Plain polarized light. Right: Crossed polars.

A bored gastropod shell in which neomorphic calcite has replaced the original lamellar CaCO3 (possibly aragonite). Traces of the original lamellar are preserved along (brown) bands of inclusions (white arrow); the neomorphic crystal boundaries cut across these structures (yellow arrows). However, also note the tendency for crystal elongation parallel to the relict lamellar textures – an indication that recrystallization was partly controlled by these earlier fabrics. Where the relict lamellae are not preserved, the calcite crystals assume a more equant, sparry habit. The boundaries of all neomorphic calcite crystals are irregular and embayed (e.g., red arrow).
The shell borings are filled with mud and silt-sized quartz and lithic grains, derived from the host sediment. Jurassic Bowser Basin. Left: Plain polarized light. Right: Crossed polars.

Neomorphism doesn’t just occur in carbonate lithologies. Here is thoroughly neomorphosed calcite in an Early Proterozoic pumice-ash tuff, Flaherty Fm. Note the highly irregular intercrystal boundaries in the coarse textures filling the shard bubble. Left: Plain polarized light. Right: Crossed polars.

Aggrading neomorphism

Aggrading neomorphism in these ooids has overprinted most of the original concentric and radial fabrics. Neomorphism is most advanced in the cores of each ooid. Maximum crystal size is 60-80 μm in ooid cores (yellow arrow), becoming finer towards the outer rims. The outer rim of each ooid has a fuzzy appearance due to recrystallization of the outer carbonate layer and the intergranular calcite cement. Several ooids have grown around fine sand-sized quartz grains (white arrow) or lime mud grains (red arrow). Left: Plain polarized light. Right: Crossed polars.

Aggrading neomorphism in this late Proterozoic cryptalgal laminated dolostone has produced a good example of structure grumeleuse, where clots of remnant micrite appear to ‘float’ within coarser, sparry, neomorphic dolomite. The boundaries of each clot are indistinct. The gradation from micrite to coarser spar may be gradual or relatively abrupt. There is a complete range of crystal sizes from micrite to 200 μm. There is a small patch of chert at lower right (arrow). Tindir Group, Alaska. Left: Plain polarized light. Right: Crossed polars.

Typical clotted texture in neomorphosed micritic dolostone. Remnants of the original cryptalgal layering are still visible. The central laminate fragment occurs with other rip-ups and disrupted, overturned microbial mats. Although less obvious, the clotted texture extends into the surrounding field of (slightly sandy) dolomite spar. Most of the crystal boundaries are irregular and embayed, indicating that neomorphism was pervasive. There are hints of relict isopachous cements on the cryptalgal fragment rims (arrows) but I cannot determine their original composition based on optical microscopy alone. Tindir Group, Alaska. Left: Plain polarized light. Right: Crossed polars.

Acknowledgement

Thanks to Annette Rogers and Kirsty Vincent, Geology at Waikato University, for access to the petrographic microscope.

Ignimbrites in outcrop and thin section

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Multiple flow units in a Mid Miocene welded ignimbrite, Chilean Altiplano.

Ignimbrite is a rock name. Ignimbrites form primarily from the collapse of eruption columns during explosive volcanism and are included in the general category of ground-hugging pyroclastic density currents (PDC). Our collective observations of PDCs over the last few centuries shows them forming primarily from explosive vulcanian and plinian eruptions that, on a human scale, have proven to be violent and destructive – witness the eruption of Mt. Vesuvius and burial of Pompeii in 79 AD (Giacomelli et al., 2003; PDF), Mt. Pelée and the destruction of Saint-Pierre in 1902 (Gueugneau et al, 2020; open access), and Pinatubo in 1991 (PDCs, airfall, and lahars).

And yet, as destructive as these recent eruptions were, they are minuscule compared to many ancient examples of thick, extensive ignimbrites that must have formed from supereruptions of cataclysmic proportions – some not so ancient; the Taupo supereruption was a mere 1800 years ago, the Oruanui (Taupo) event 26,500 years ago (Wilson et al, 2006). The shear power and violence of this type of eruption is difficult to comprehend.

The original definition of ignimbrite by P. Marshall (1932, 1935 paper reproduced here), based on a comparison of New Zealand examples with past events in the Valley of Ten Thousand Smokes, incorporated field-based criteria such as thickness, relationship to topography (do the deposits thicken in valleys?), consist primarily of juvenile pumice and ash, and the presence of welding. For many years welding was considered a prerequisite for ignimbrite status.

The decades following Marshall’s inspired investigations saw the term ignimbrite extended to all manner of pyroclastic flows – thick, thin, laterally extensive or of more local distribution, hot or cold including those that are not welded (e.g., the 26.5 thousand year old Oruanui ignimbrite, Taupo, Wilson et al., 2006; PDF). Giordano and Cas (2021) have attempted to rationalize this general terminological conflation with the following definition: an ignimbrite is “…the rock or deposit formed from pumice and ash- through to scoria and ash-rich pyroclastic density currents” regardless of thickness, areal extent, volume, composition, crystal content, relationship with topography, or temperature. This sounds like an anything-goes kind of definition, but Giordano and Cas also develop an ignimbrite classification scheme that distinguishes those formed from single vent eruptions (vulcanian and plinian, like Vesuvius and Pinatubo), from those of much greater magnitude and intensity that form during caldera collapse (issuing from multiple vents/fissures that ring the caldera); their diagram is reproduced below. This categorization is based on measurable quantities such as thickness, areal extent (flow run-out), volume, and aspect ratio. Dispersal area is controlled by the mass flow rate during an eruption and is a proxy for eruption intensity; likewise, flow volume is controlled by flow rate and the duration of the eruption and is a proxy for eruption magnitude.

Classification of ignimbrites formed from single vent eruptions, and the spectrum of multiple vents or fissures associated with caldera collapse, of increasing eruption intensity and magnitude. Redrawn from Giordano and Cas, 2021.

[Aspect ratio is calculated as the average thickness versus the average lateral extent. The concept was introduced to volcanology by George Walker (1983) to distinguish the conditions that create ignimbrites (explosivity, eruption magnitude and intensity)]

[Eruption intensity refers to the rate at which the pyroclastic mass is ejected]

[Eruption magnitude refers to the total mass erupted – the scale of erupted volumes is also encapsulated in the Volcanic Explosivity Index, or VEI]

Common ignimbrite traits

The following list is based primarily on field and petrographic characteristics. In addition to the references cited, I have also drawn on the chapter on ignimbrites by Freundt et al., 2000 (Encyclopedia of Volcanology).

Deposit thicknesses range from decimetres to 200 m and more. Thickness can be variable even within single flow units.
Their distribution is largely determined by topography, thickening in topographic lows and thinning over ridges. Their geometry reflects the dominant control of gravity from collapsing eruption columns. Individual flows are capable of blanketing landscapes for 10s of 1000s km².
Calculation of the volume erupted needs to account for magma fragmentation, loss of volatiles, and the amount of fine ash dispersed far beyond the observed area of dispersal. It also needs to account for differences in pumice and glassy ash density compared with the equivalent dense, non-fragmented rock (i.e., solid lava). Therefore, erupted volumes are commonly quoted in terms of their dense-rock equivalent (DRE) – i.e., the equivalent volume of non-fragmented lava that in most cases would be rhyolite or dacite.
Thick ignimbrites commonly appear massive, but on close inspection may reveal several stacked flow or cooling units separated by co-ignimbrite airfall tephra, thin pyroclastic surge deposits, or abrupt changes in grain size or crystal concentrations and composition.
Grain size grading and stratification may develop in the upper and distal parts of flow units, reflecting waning turbulence in the former, and the transition to laminar fluid flow in the latter.
Most ignimbrites have rhyolite through dacite compositions. They have an abundance of juvenile pumice, crystals, and pyroclasts of all sizes, although fine ash to coarse lapilli are more common. The term juvenile refers to derivation from the erupted magma.
Within single flow units, crystal and pyroclast compositions may vary vertically and laterally. Some ignimbrites are crystal-phenocryst rich, others are crystal-poor. Such variations reflect the availability of materials in the magma, particularly as the magma chamber empties. Very large eruptions (supereruptions) may tap more than one magma chamber, and the magma composition in each may vary. This variation is reflected in the compositional zonation within and among closely spaced flow units, although mechanical winnowing of some crystal phases may also take place in the eruption column and ensuing PDC (see Wilson et al, 2021 for an excellent review of magma chamber evolution).
Phenocrysts are commonly broken by the intensity of the eruption and by mechanical diminution during turbulent pyroclastic flow.
Country rock derived from vent or fissure walls, or plucked from the substrate beneath a PDC may be added to the volcaniclastic mix.
Welding is regarded as a syndepositional process. Welding, (also referred to as sintering) of pumice clasts and glass shards occurs when an ignimbrite is hot enough for glass to maintain a relatively fluid or ductile state during deposition. Experimental data on glass rheology indicates that temperatures greater than 800^oC are required for this to take place (Lavallee et al. 2015; open access). Welding results in compression and stretching of pumice clasts, producing flame-like structures, or fiamme. Welding of finer ash produces agglutinated glass shards that, at a microscopic scale, appear to flow around phenocrysts and larger pyroclasts. Welding may be sufficiently pervasive to produce crude banding or layering (also called eutaxitic texture) that resembles flow-banding in felsic lava flows. Note that this kind of layering is not the same as that produced by laminar fluid flow in some PDCs.
Eutaxitic textures, a term usually reserved for welded ignimbrites, refers to the layering produced by the flattening and collapse of pumice fragments, and the agglutination of glass shards during sintering.
Columnar cooling joints are common in thick ignimbrites.
Porous, non-welded ignimbrites are susceptible to fumarolic alteration of glass and phenocrysts during cooling.

Ignimbrites in outcrop

A panorama of the 1.21 Ma Ongatiti rhyolite-rhyodacite ignimbrite (Waikato, New Zealand), at least 11 m thick (base not exposed). At this locality (Castle Rock) there is no indication of extensive welding. At other localities, this ignimbrite is welded and contains several flow units ( Briggs et al., 1993) – see the thin section examples of welding in this ignimbrite. Elsewhere the ignimbrite thickens to 35 m and more. The eruption centre was the Mangakino Caldera (Taupo Volcanic Zone), about 20 km south of these cliffs. The dense-rock equivalent volume (DRE) is more than 500 km³.

Typical weathered, pockmarked exposure of the Ongatiti Ignimbrite. The ‘holes’ are formed where soft pumice clasts are removed by weathering – giving us a good visual indication of their abundance. The largest clasts here have 40 mm long axes. Clast size sorting is extremely poor. Ignimbrite bluffs in this area are popular with rock climbers.

Janine Krippner is intrigued by this pockmarked exposure of the Ongatiti Ignimbrite. The ‘holes’ are formed where soft pumice clasts are removed by weathering – giving us a good visual indication of their abundance. The largest clasts here have 40 mm long axes. Clast size sorting is extremely poor. Ignimbrite bluffs in this area are popular with rock climbers.

A relatively intact, non-sintered pumice clast with preserved vesicles (there is some iron oxide alteration), Ongatiti Ignimbrite (location as above).

Non- to weakly welded, fine lapilli-sized pumice and rhyolite pyroclasts dispersed in much finer grained glassy ash. Note the variation in shape and angularity, and extremely poor sorting. Part of the Okataina Caldera (about 320 ka) exposed along the shore and road cut of Lake Rotoma. Taupo Volcanic Zone, New Zealand. Coin is 23 mm diameter.

Non- to weakly welded, fine pumice and rhyolite lapilli dispersed in much finer grained glassy ash. Note the variation in shape and angularity, and extremely poor sorting. Part of the Okataina Caldera (about 320 ka) exposed along the shore and road cut of Lake Rotoma. Taupo Volcanic Zone, New Zealand. Coin is 23 mm diameter.

Welded ignimbrite exposed along the faulted shoreline margin of Lake Rotoma, part of the Okataina Volcanic Centre (about 320 ka). Most pumice fragments are stretched and flattened (arrows) but a few denser varieties are unchanged (yellow arrow). The ignimbrites are intercolated with flow banded rhyolites, spherulitic obsidian, and coarse breccias. Taupo Volcanic Zone, New Zealand. Bar scale 30 mm.

Intensely welded, flattened, and stretched pumice fragments forming typical fiamme. Welding is pervasive in this ignimbrite, part of the Late Pliocene Coroglen Subgroup, exposed north of Whiritoa, Coromandel Volcanic Zone.

The Mamaku Ignimbrite (240,000 yr) was one of the main eruptive products during the Late Pleistocene caldera collapse that formed Lake Rotorua. The upper, non-welded portion of the PDC was subjected to fumarolic discharge during cooling that resulted in pockets of cement hardened deposits. Subsequent erosion removed the soft, non-cemented ignimbrite, leaving upstanding remnants of the hardened rock preserved as steep-sided mounds capped by rocky spires and blocks. These landforms are called Tors, or Inselbergs. The top of the Mamaku Ignimbrite is littered with tors – shown in this Google Earth image. Taupo Volcanic Zone, New Zealand.

A tor at the top of the Mamaku Ignimbrite (west of Rotorua). The mound and spire are about 9m high. The general relief here also indicates the amount of denudation since ignimbrite emplacement. Taupo Volcanic Zone, New Zealand.

Road cut through a Mamaku Ignimbrite tor (Hwy 5, west of Rotorua). The eroded top of the Ignimbrite is draped by younger airfall tephra. See Google Earth image above for location. Taupo Volcanic Zone, New Zealand.

Eroded remnants of a Middle Miocene welded ignimbrite bordering Salar La Grande, Chilean Altiplano (orange hues, arrow top left). The high ground behind the ignimbrite is mostly composite andesitic and basaltic cones of similar age. The eruptive episodes were associated with Mid-Miocene caldera collapse. The ignimbrite is lapped by more recent, arid, alluvial fans, the alluvium derived from Mid Miocene basalts and andesites. The Salar is capped by thick (white) gypsum and halite crusts.

The ignimbrite (same unit as image above) is composite, consisting of overlapping flow units (white arrows) separated by subtle changes in pumice and crystal content, the contacts weathering out as apparent horizontal joints. About 5 m of section in this view. In the background, a nicely preserved Middle Miocene volcanic cone, plus a peekaboo view of Salar La Grande.

A closer look at the welded, flattened pumice, accentuated by weathering (same unit as image above). There is a crude alignment of the flattened clasts. Hammer lower left.

Ignimbrites in thin section

Welded 1.21 Ma Ongatiti Ignimbrite (located in a quarry at Hinuera). Left: Broken plagioclase phenocrysts are surrounded by flattened and stretched glass shards – compression has produced flow patterns around the crystals. The remnants of bubble-wall textures are seen in some shards (arrows). Plain polarized light. Right: The glass component is isotropic under crossed polars.

Welded Ongatiti Ignimbrite (same location as above). Prominent flow textures and flattening of glass shards occurred while the glass temperature was high enough to sustain ductile behaviour. The ragged phenocrysts are plagioclase, broken during eruption or during turbulent transport in the PDC. Plain polarized light.

Intensely welded Matahina Ignimbrite (322 ka), part of the Okataina Volcanic centre. Taupo Volcanic Zone, New Zealand. The glass appears fuzzy (in PPL) because it has been altered by fumarolic activity during the post-depositional cooling phase. The phenocrysts are a mix of broken and complete crystals. Left: Plain polarized light. Right: Crossed polars. Note the preserved crystal faces on the plagioclase top left (p). Of the two broken plagioclase fragments, one is zoned (z), the other untwinned (it has good cleavage). There is a single, small pyroxene (px). The squiggly crystal at bottom centre may be a resorbed plagioclase (r): it has cleavage, but no twinning. Resorption occurs when there is disequilibrium between the magma melt and the solid crystal phase.

Pumice fragment showing good preservation of vesicles. Plain polarized light. Taupo Volcanic Zone, New Zealand.

$A seriously sintered, flattened pumice fragment, with collapsed vesicles. The fractured quartz phenocryst probably indented the pumice fragment during compression while the glass was still relatively fluid. Individual glass shards in the surrounding ash have also been welded together forming pronounced eutaxitic texture. Waiotapu Ignimbrite (710 ka). Taupo Volcanic Zone, New Zealand. Plain polarized light.$

A seriously sintered, flattened pumice fragment, with collapsed vesicles. The fractured quartz phenocryst probably indented the pumice fragment during compression while the glass was still relatively fluid. Individual glass shards in the surrounding ash have also been welded together forming pronounced eutaxitic texture. Waiotapu Ignimbrite (710 ka). Taupo Volcanic Zone, New Zealand. Plain polarized light.

The rheology of volcanic glass during intense sintering is nicely illustrated here with flow and compression around a quartz phenocryst (lower centre). Most of the glass shards have been flattened. The glass in this and the previous image has been devitrified. Waiotapu Ignimbrite (710 ka). Taupo Volcanic Zone, New Zealand. Plain polarized light.

Fresh pumice from the Taupo supereruption, about 1800 years ago. No welding or compaction. The fragment is basically a collection of vesicles held together by a glass lacework. Plain polarized light. Taupo Volcanic Zone, New Zealand.

Pumice fragments from a thick PDC in the Paleoproterozoic Flaherty Fm. Belcher Islands. The associated volcanics (lava flows and pillow lavas) are basaltic and at least partly submarine. There is no indication of sintering. One clast has elongated vesicles and small, black, opaque hematite crystals (top left, arrow). Vesicles are filled with chlorite (yellow arrow) and calcite; the main cement is calcite. There are a few grains of altered basalt (red arrow). The original glass has been replaced by Fe oxides. Plain polarized light.

Acknowledgement

Many thanks to Kirsty Vincent, Earth Sciences, Waikato University for access to the petrographic microscope.

Accretionary aggregates and accretionary lapilli

Block and ash flows

Ignimbrites in outcrop and thin section

Volcanics in outcrop: Pyroclastic density currents

Volcanics in outcrop: Secondary volcaniclastics

Volcanics in outcrop: Lava flows

Volcanics in outcrop: Pyroclastic fall deposits

Fluid flow: Froude and Reynolds numbers

Sediment transport: Bedload and suspension load

The hydraulics of sedimentation: Flow regime

Glauconite in thin section

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Glauconite filling of rotaliid foram chambers

Thin section of glauconite filling rotaliid foram chambers

Glauconite – an important stratigraphic and paleoenvironmental indicator

Glauconite is one of the most recognizable minerals in the entire sediment suite, primarily because of its green hues, its common peloidal grain morphology, and microcrystalline structure. Green-ness does apply to a few detrital minerals like amphiboles, pyroxenes and chlorite, but they can be distinguished by their crystal habit and prominent cleavage; sedimentary glauconite is always microcrystalline.

Glauconite is a sheet silicate, distinctive in its high concentrations of Fe²⁺, Fe³⁺, and K. It occurs as different types depending on the relative proportions of Fe and K. Well-ordered (mature) glauconite is high in K (8% to 9%, derived from sea water) and its crystal structure is closest to that of micas. Precipitates with lower K contents tend to have more clay-like structure, particularly smectite (a mixed crystallographic bag of expandable clays). Total Fe is commonly >15%, and the amount of Fe³⁺ is usually > Fe²⁺ (B. Velde, 2014). Glauconites with low K and high Fe tend to have brown-yellow hues. Most glauconites begin precipitation as K-poor smectites that, over about 10⁵ to 10⁶ years, become increasingly K-rich and crystallographically ordered, or mica-like (Odin and Matter, 1981). The increase in crystallographic ordering can be measured using X-ray diffraction.

Glauconite is a marine, authigenic mineral that precipitates at very shallow depths beneath the sediment-water interface. This requires very low sedimentation rates, normal salinity and pH of the ambient seawater, and suboxic conditions where both Fe³⁺ and Fe²⁺ are thermodynamically stable. The oxidation of organic matter in peloids (particularly fecal pellets) and invertebrates may help push the oxidation state of Fe towards the 2+ state. Note that soluble Fe²⁺ will oxidize rapidly to relatively insoluble Fe³⁺ if exposed to normal, oxygenated, shelf sea water, forming compounds like limonite and goethite. Modern glauconite commonly accumulates at depths <300-500m on siliciclastic and carbonate platforms, shelves, and continental slopes. Ancient accumulations are frequently encountered as greensands.

The most common granular form of glauconite is sand- and silt-sized peloids. Other grain morphologies include vermiform (worm-like), tabular, mammillated, and composite types (e.g. Triplehorn, 1966). Glauconite can also replace or infill the chambers of bioclasts like gastropods, forams, bryozoa, corals, and the perforations in echinoid plates and spines. In hand specimen, they usually exhibit as dark green to black, sub-millimetre, ovoid pellets.

Accumulations of glauconite in the rock record are of great value because:

They indicate a unique set of paleoenvironmental conditions.
They are excellent indicators of stratigraphic condensation during transgression (look for a maximum flood surface at the top of the condensed section).
They may coexist with phosphate nodules and phosphatized fossils (that also indicate condensed stratigraphy).
Although they form authigenically, glauconite can be reworked locally. Glauconite can also be moved to deeper waters by sediment gravity flows. It is soft, friable, and has a hardness of 2 on the Moh scale, matching that of gypsum. Thus, it rarely survives beyond the first cycle of sedimentation.

Glauconite and glaucony

The name glaucony was introduced by Odin and Letolle (1980) for all forms of glauconite characterized by smectite clay crystal structure; the term glauconite was reserved for the mica-structured mineral end-member. However, the name ‘glauconite’ is firmly entrenched in decades of literature and for better or worse, I have continued with this literary laziness.

Glauconite (glaucony) in thin section

Chatham Rise

We begin with some glauconite-rich deposits from Chatham Rise, a ridge that is part of the submarine extension of Zelandia continent. The deposits form part of a highly condensed stratigraphic package dating back to the Oligocene and have been interpreted as palimpsest (relict) remnants of glacially lowered sea levels. Glauconite peloids that presently carpet the sea floor atop Chatham Rise have been K-Ar dated at 5-6 million years old (Nelson et al., 2021).

$Two samples from Chatham Rise. The dominant morphology of glauconite is peloidal, with a few composite aggregates and some fossil infills (mainly benthic and planktic forams). The range of colours, from yellow-brown to grass green reflects variation in their chemical maturity – darker greens are higher in K (an average 7.2% K) and have greater crystallographic ordering determined from X-ray diffractometry. Those tending towards more brown hues have higher Fe2O3. Many peloids have a brown rim of limonite (Fe2O3). Individual peloids may also show internal colour gradation from lighter green or brown exteriors to darker green interiors. Both images plain polarized light. Images courtesy of Cam Nelson. https://www.geological-digressions.com/contributors/contributions-from-cam-nelson/ -Peloids are ovoid to tabular with smooth surfaces. - Most peloids have quasi-radial or interconnected networks of cracks (dark brown – Fe oxide?). This is a common feature that mechanically weakens the grains such that they are unlikely to survive vigorous reworking or abrasion. - Both examples contain planktic and benthic forams that show the beginnings of Fe-oxide-glauconite chamber linings. - Image left contains fragmented echinoderm plates and spine cross-sections where perforations are partly filled with red-brown Fe oxide-glauconite.$

Two samples from Chatham Rise. The dominant morphology of glauconite is peloidal, with a few composite aggregates and some fossil infills (mainly benthic and planktic forams). The range of colours, from yellow-brown to grass green reflects variation in their chemical maturity – darker greens are higher in K (an average 7.2% K) and have greater crystallographic ordering determined from X-ray diffractometry. Those tending towards more brown hues have higher Fe2O3. Many peloids have a brown rim of limonite (Fe2O3). Individual peloids may also show internal colour gradation from lighter green or brown exteriors to darker green interiors. Both images plain polarized light. Images courtesy of Cam Nelson.
– Peloids are ovoid to tabular with smooth surfaces.
– Most peloids have quasi-radial or interconnected networks of cracks (dark brown – Fe oxide?). This is a common feature that mechanically weakens the grains such that they are unlikely to survive vigorous reworking or abrasion.
– Both examples contain planktic and benthic forams that show the beginnings of Fe-oxide-glauconite chamber linings.
– Image left contains fragmented echinoderm plates and spine cross-sections where perforations are partly filled with red-brown Fe oxide-glauconite.

Modern sea floor, Three Kings Islands, northern New Zealand

The shallow, modern sea floor around Three Kings Islands is accumulating bioclastic sediment, dominated by bryozoa, forams, echinoids, barnacles, and molluscs. In this mix are glauconite pellets and glauconite infilled bioclasts.

Left: An irregular peloid with cracked, vermiform glauconite. The dark brown-black, opaque (and isotropic) material, that along its edges is yellow-brown, may be limonite. Right: Peloidal glauconite in shades of green-brown and yellow-brown. The cracks typically taper towards the peloid centre (arrows). The cracks may have formed by syneresis (a process of fluid expulsion typically associated with the transformation of gels to crystalline materials – I think the jury is still out on this interpretation). Cracks will render a peloid more vulnerable to disintegration from vigorous reworking; thus, the present examples have probably moved very little. Both images in plain polarized light. The bar scale applies to both images.

The microcrystalline habit of glauconite presents as a speckled appearance, best seen under crossed polars. The three grains (arrows) are the same as in the plain polarized image above. Grains with first-order grey interference colours are quartz. The bioclasts are probably molluscs and barnacles.

Left: Three things to note here: 1. The planispiral foram chambers have incipient glauconite fill, that along the inner margin of the test wall are transforming to pale green (red arrow). The radial crystallite wall structure is still visible. 2. Three ovoid glauconite peloids to the left of the foram display variations in colour that probably correspond with a range of maturities. The larger grain has developed a few cracks (black arrow). 3. The grain top left resembles barnacle plications; here too glauconite has begun to form with the outer grain rim developing green hues (arrow). Plain polarized light.
Right: A rotaliid foram chamber (radial wall structure) partly filled by yellow-green glauconite and a few silt-sized fragments of quartz. The left side of the fill shows mammillated morphology. Plain polarized light.

An oblique section through an echinoderm spine where the internal perforations are filled with immature, brownish glauconite. Left – plain polarized light. Right – crossed polars. The bar scale applies to both images.

Glauconite infill of fossils

Examples from Oligocene bioclastic limestones, Te Kuiti Group, New Zealand

The emphasis in this image pair is glauconite filling invertebrate microporosity: on the left the perforations in an echinoderm plate (e); at centre-right a longitudinal section through bryozoa zooids (b). An irregular, composite mass of glauconite has also formed along the echinoid plate margins also contains silt-sized quartz grains. The bioclasts are cemented by coarse calcite spar. Left – plain polarized light; Right – crossed polars.

The outer chambers of this rotaliid foram are filled with yellow-green glauconite; the inner chambers with calcite spar. The radial wall structure is still visible although much of the calcite has been replaced by limonite and glauconite. The outer margins of the foram may have an isopachous calcite rind (arrows). The remainder of the cements are coarse calcite spar. Plain polarized light.

Glauconitic micaceous arenite

– Glauconite peloids up to 0.5mm across (10%). Some glauconite grains are deformed around stronger quartz and feldspar grains (by compaction).
– Quartz grains (70-75%) are mostly monocrystalline (there are a few polycrystalline grains) and show a mix of unstrained and strained extinction. Grains are angular to well rounded, but some of the angularity is due to calcite replacement of quartz (some examples shown by white arrows).
– Untwinned K-feldspars (5-10%) show varying degrees of sericite alteration. Some appear superficially like lithic grains.
– Lithics (I) (~5-8%). Be careful to distinguish actual lithics from altered feldspar grains.
– White micas (1-2% – red arrows) show some alteration to chlorite (bluish interference colours).
– The framework clasts and micas have a crude alignment along the top left/bottom right diagonal.
– Cement is almost entirely coarse calcite. Contact between the calcite cement and framework grains is irregular, indicating silica replacement.

Glauconitic arenite

$The sandstone framework here is compositionally similar to the arenite above, but note the changes in texture and fabric: -The grains are more densely packed, with significantly less calcite cement and decreased calcite replacement of framework grains. - The original peloidal shape of the glauconite grains has been distorted by compaction around stronger quartz-feldspar and squeezed between some grains (red arrows for some examples). - Compaction has fractured some quartz grains – fractures are filled by calcite and possibly clays.$

The sandstone framework here is compositionally similar to the arenite above, but note the changes in texture and fabric:
– The grains are more densely packed, with significantly less calcite cement and decreased calcite replacement of framework grains.
– The original peloidal shape of the glauconite grains has been distorted by compaction around stronger quartz-feldspar and squeezed between some grains (red arrows for some examples).
– Compaction has fractured some quartz grains – fractures are filled by calcite and possibly clays.

Carbonates in thin section: Forams and sponges

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A couple of rotaliids in various states of comminution (possibly Amphistegina – the one labelled ‘r’ is broken along its left margin), and an evolute planispiral foram (centre) in bioclastic grainstone that also contains plenty of bryozoa (b), echinoid plates (e) and spines, and barnacles (ba). The central foram is partly filled with glauconite. Oligocene, New Zealand. Plain polarized light.

Two important contributors to bioclastic limestones – foram tests and skeletal sponge spicules

Foraminifera

Foraminifera (forams) are single celled, marine to brackish water organisms that inhabit the photic zone (their diet consists of algae, diatoms, and other photic zone micro-organisms). They construct exoskeletons (tests) either by secreting calcium carbonate, or by gluing together silt and clay-sized sediment particles (agglutinated forms); there is also a small group of chitinous forms. They are divided into two broad groups: planktic species that spend their lives within the water column, and benthic forms that live on or in the sea floor sediment. A few species encrust other larger shells, corals, and bedrock. Benthic forams are by far the largest and most diverse group.

Forams are sensitive to changes in environmental conditions such as substrate, water depth, salinity, and temperature and thus are excellent paleoenvironmental and paleobathymetric indicators. Their sensitivity to temperature is also reflected in the partitioning of stable oxygen and carbon isotopes, providing excellent proxies for oceanic paleotemperatures. They are also one of the most important fossil groups for high resolution biostratigraphy, particularly in Cretaceous and Cenozoic rocks, a development for which the oil exploration industry can claim much credit.

Forams are microscopic, mostly less than a millimetre across, except for a few species like the iconic Nummulites that in some cases can exceed 50mm diameter. You will need a decent binocular microscope to get any detail out of them; modern foram studies routinely use a scanning electron microscope for species ID.

The variety of test form ranges from single chambered species to those with multiple chambers arranged in linear and flat spiral arrays, and conical spiral forms that look superficially like small gastropods. Some of the more common forms are shown below.

Some of the more common foram test shapes. The original figure is from Loeblich and Tappan, 1964 (the current figure was accessed and modified from https://www.ucl.ac.uk/GeolSci/micropal/foram.html )

Foram classification is based on test morphology, wall structure, and aperture geometry.

The test morphology one sees in thin section depends on the orientation of the fossil. In this situation, the structure of the test wall can be an important guide, at least to the taxonomic level of order.

There are four main orders

Textulariids: Agglutinate species grow their tests by ‘gluing’ together small sedimentary particles with calcite or silica cement. In thin section they typically appear dark, almost isotropic, although silt-sized quartz, feldspar, and fossil fragments are usually visible.

A modern agglutinate foram (Textulariid) from northern New Zealand made up of clay and silt-sized sediment particles (zoom in to see these). Left plain polarized light. Right A schematic representation of a Textulariid chamber and agglutinate particles.

Rotalìids (Hyalina – hyaline means almost transparent). Rotaliids secrete either calcite or aragonite (never a mix of the two) in crystallites arranged in radial fashion with c-axes normal to the test wall. A new layer of radial crystals is precipitated for each chamber addition. Thus, in thin section the foram test walls appear layered. Under crossed polars there may be a pseudo-extinction cross that will rotate with the microscope stage.

This group includes benthic and planktic species, including well-known genera like Nummulites and Globigerina. Test walls have fine perforations (perforate) – these are easy to see in planktic forms. Some species contain microgranular calcite, and a few have walls constructed of large, perforated crystals.

An Oligocene Amphistegina (a Rotaliid) in cross-section showing the characteristic layered and radial, almost prismatic crystal structure. Crystallites are generally normal to the test wall. Plain polarized light. A simplified schematic of this texture is shown on the right. Under crossed polars, crystal extinction appears to sweep across the foram section.

Several Rotaliids, probably Amphistegina, in an Oligocene bryozoa (b)-foram bioclastic grainstone. Chamber walls are layered with crystals arranged in characteristic radial patterns. Cements are mainly drusy calcite. Pale green patches are incipient glauconite cements. Plain polarized light.

Details of some Recent Rotaliid forams showing the characteristic layered and radial-prismatic crystal textures. Left plain polarized light. Right crossed polars. Note the pseudo-extinction crosses in the forams; each cross will rotate in concert with the microscopic stage. The echinoderm spine cross-section (left) is a single crystal of calcite.

SEM image of the iconic planktic foram, Globigerina, a trochispiral Rotaliid. The test is peppered with micron-sized perforations, or pores. Image credit: Bolli, H. M., Saunders, K. Perch-Nielsen, 1985

SEM image of the iconic planktic foram, Globigerina, a trochospiral Rotaliid. The test is peppered with micron-sized perforations, or pores. Image credit: Bolli, H. M., Saunders, K. Perch-Nielsen, 1985

A mudrock with abundant planktic forams. Zoom in on the large foram lower left to see the radially arranged wall crystallites and spinose ornamentation. Most of the forams are intact, indicating relatively gentle suspension deposition from the overlying water column. There is a single fragment of calcareous algae bottom right. The star-shaped fossil (top right) looks similar to modern star-shaped forams (Baclogypsina sphaerulata) found on some Indo-Pacific beaches, particularly Okinawa where they are known as star sands. Sample source and age unknown.

Foraminifera ooze containing abundant planktic rotaliid forms (Globigerina-like), plus a few miliolid forams (microgranular texture – see description below) – note the biserial form near bottom left (b). The rotaliids show typical radial crystal walls with perforations extending the entire wall thickness, and spinose ornamentation, shown in greater detail in the inset image. Most chambers are filled by coarse crystalline calcite. Image courtesy of Mark Lawrence, GNS.

Fusulinids are an order that died off during the end Permian mass extinction. The walls of this group consist of microscopic, equigranular calcite. Many, but not all, are perforate.

Left: Bedding of Permian grainstone exposing a Fusulinid multitude; the elongate forams have been aligned by marine currents. Right is a schematic of typical fusuline microgranular calcite wall structure. The c-axes of individual calcite crystals are oriented randomly – under crossed polars the crystallites will also extinguish randomly as the microscope stage is rotated. The example is from South Bay, Ellesmere Island.

Miliolids Porcellanous (smooth and commonly shiny) – This group has walls constructed of high-magnesium, acicular to equant calcite crystals, generally lacking preferred orientation. Tests are imperforate (cf. the Fusulinids).

A biserial foram with Miliolid, or porcellaneous wall structure (left), consisting of needle and/or equant calcite crystallites that have no preferred orientation. Right is a schematic representation of this wall structure.

Sponges

A nice collection of extant sponges living amongst Porites corals in Columbia reef, Cozumel. On the right, yellow tube and encrusting species; lower centre-left pale blue vase sponges; and lower left and top, three large barrel sponges. Image courtesy of Charlie Kerans.

Sponges (phylum Porifera) are the simplest of multicell invertebrates. They survived multiple extinction events since the Cambrian; there are hints that they may have evolved during the late Precambrian. Most forms have an internal scaffolding-like skeleton made up of solid rod-like structures called spicules (or sclerites), that are bound together to form a semi-rigid framework, around which the spongin grows and is structurally supported (spongin, as in a bath sponge, is an insoluble protein). The spicules are composed of amorphous silica (opal-like), or high-magnesium calcite – the former are more common. Siliceous sponges are also called glass sponges. Spicules also focus light in much the same way as optic fibres – an interesting discovery. The spicules disaggregate when the animal dies. Sponge taxonomy is based primarily on spicule geometry.

Sponge spicules come in many shapes, usually less than a few millimetres long. They occur as single rays, pointy at the growing end, or as complex architectures that radiate from a central point. Each species has a unique spicule size and shape. Calcareous spicules (Calcarea) are single crystals. Siliceous spicules have a central canal.

All spicules are prone to recrystallization during burial diagenesis, partly because of their composition (the amorphous silica will recrystallize to quartz), but also because of their relatively high surface area. Siliceous spicules may also be replaced by carbonate and in the process provide an important source of dissolved silica to ambient geofluids.

Disaggregated siliceous sponge spicules at different orientations in Recent sediment, northernmost New Zealand. The longest is about 2.2mm. The spicules in this group are called Tetraxons having three rays of equal length plus a much longer ray, all radiating from a single point. Variations on this theme include spicules with one, three, and six or more rays. The central canals appear as dark lines within each ray.

A grain-mount thin section of Recent bioclastic sediment from northern New Zealand. There are abundant Rotaliid forams (radial wall structure) plus a variety of intact and broken sponge spicules (s) that present as clear, colourless, pin-like rods. Plain polarized light.

Acknowledgement

Many thanks to Kirsty Vincent, Earth Sciences, Waikato University for access to the petrographic microscope.

Carbonates in thin section: Bryozoans

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Devonian crinoid stems and disarticulated columnals in the bioclastic mix with fragmented fenestrate (f) and encrusting bryozoans (e).

Bryozoa – are a complex but important group of reef and non-reef invertebrates and contributors to bioclastic sediment.

Bryozoa are an incredibly diverse group of metazoans with a geological record extending from Early Ordovician to Recent. There are more than 6000 extant species and probably 2-3 times that number of fossil species. They were once classified as plants, or moss animals, because in many cases that’s exactly what they look like. But they are metazoans, tiny as individuals (zooids, commonly less than 1 mm diameter), but as colonial organisms can construct sizeable, rigid mounds and thickets. The majority are marine; a small group, the order Ctenostomata are non-marine. Bryozoa are important contributors to bioclastic sediment in shallow, cool-temperate water depositional systems. The larger buildups also provide an ecological niche for other organisms.

The zooids are housed within structures built of calcium carbonate, chitin, or a mix of both; the solid skeletal structure of a colony is called a zoarium (pl. zoaria). The exoskeleton constructed by an individual zooid is called a zoecium (pl. zoecia). Most species secrete either calcite or aragonite exoskeletons. Only the carbonate forms have good preservation potential (so it’s worth keeping in mind the likelihood that in fossil bryozoan assemblages there will probably be “missing” species). Many species contain a mix of calcite and aragonite. For modern bryozoans the proportion that secretes aragonite is greatest in tropical realms, generally decreasing in cooler waters at higher latitudes. The Mg content of calcites is also highly variable, ranging from 0 to 16 (mol)% (Smith, 2014 – for a review of some modern bryozoans).

Common forms range from low relief encrusters on bedrock, shells, and seaweed, to rigid frameworks and thickets that are delicately or robustly branched (some coral-like), palmate (shaped like a palm tree), or fenestrate (window-like) and lacy, anastomosing sheets, massive domes, and even corkscrew structures commonly labelled as Archimedes screws. Colonies a few centimetres long may house a million and more zooids.

A few of the common bryozoa forms. Each form has a 3-dimensional aspect that can potentially create large (decimetre-scale), complex structures. Individual skeletal cavities – zoecium – are indicated. In fenestrate forms, the (relatively) large openings in the zoarium do not accommodate the living animal; however, they can be filled with sediment and cements as is the case for the zoecia (see the images below).

Secretion of calcium carbonate confers a degree of rigidity to bryozoa, but in the grand scheme of sediment production and dispersal, they are relatively delicate structures. This is particularly the case for delicately branched and sheet or plate-like zoaria. As fragments in grainstones and rudstones, their identification even to the level of bryozoan order can be difficult. In thin section, fragment identifications generally rely on a relatively ordered arrangement of zoecium that are regularly spaced and of similar size and shape. Note that the size and shape of the zoecium may change as the orientation of the fragment changes.

Modern bryozoa encrusting a gastropod. The zooids are arranged in a brick-like pattern. Hahei, NZ.

Modern rigid, calcareous branched bryozoa. Zoom in to see the zooids along each branch. Specimen is from Florida Keys.

Differentiation from some other phyla

Corals:

The corallites of solitary or colonial corals are significantly larger than bryozoan zoecium
Corals are composed of aragonite, that in most cases will recrystallize to low magnesium calcite during burial diagenesis.
Corallites have well defined internal structure, founded on septa (vertical plates of aragonite) that radiate from a central column – zoecia have no such internal structure.

Barnacle plates:

Basal sections of barnacle plates commonly show an array of large pores that resemble bryozoan zoaria. However, the barnacle fragment may also show its characteristic plicate layered structure. Bryozoa exoskeletons have no well-defined layering.

Echinoderms plates:

There is a superficial resemblance between fenestrate bryozoa and perforated echinoderm plates, where pores in the latter are commonly arranged linearly or in quasi-radial patterns. Echinoderm plates also consist of a single calcite crystal – this will show as uniform extinction across the entire plate. Bryozoa plates consist of a myriad, microscopic acicular crystals that during diagenesis will recrystallize to calcite spar.

Cross section through a Cheilostome bryozoan colony. Note the change in zoarium shape and size as the orientation across the sample changes from mostly longitudinal (left – plain polarized light) to transverse (right – crossed polars). The zoecia have been filled with drusy calcite cement.

An abraded fragment of modern, lacey or platy bryozoa showing variable zoecium size (average 200 μm zoecium width) and shape. Zoecia walls are constructed with fibrous aragonite crystals that, in the expanded view (right - crossed polars), are crudely aligned parallel to zoecia walls. Offshore Three Kings Islands, northern (subtropical) New Zealand, about 30 m depth.

An abraded fragment of modern, lacy or platy bryozoa showing variable zoecium size (average 200 μm zoecium width) and shape. Zoecia walls are constructed with fibrous aragonite crystals that, in the expanded view (right – crossed polars), are crudely aligned parallel to zoecia walls. Offshore Three Kings Islands, northern (subtropical) New Zealand, about 30 m depth.

Two kinds of bryozoa from recent bioclastic sediment, Three Kings Islands (northern New Zealand): Plate-like forms (p), and a single form with anastomosing zoecia walls (a) – Note the micron-sized pores in this specimen. For contrast, the solitary corallite (c) is an order of magnitude larger than the zoecia and shows the typical radial disposition of septa (partly sediment filled). Right image shows the anastomosing bryozoa in greater detail. Plain polarized light.

An Oligocene bryozoa-dominated limestone. There are numerous longitudinal ((bl) and transverse (bt) sections through the bryozoa fragments. In the bioclastic mix are a couple of foraminifera, the planispiral form (centre) partly filled with glauconite. There are also a few abraded echinoderm plates (e) that are encased in syntaxial calcite cement. Left: plain polarized light; Right: crossed polars.

Oligocene Awakino limestone dominated by bryozoa, including a specimen showing good palmate structure (p – a longitudinal view) where individual tubes splay upwards. In the bioclastic mix are a few small foraminifera and small fragmented echinoderm plates. Plain polarized light.

Acknowledgement

Many thanks to Kirsty Vincent, Earth Sciences, Waikato University for access to the petrographic microscope.

Carbonates in thin section: Echinoderms and barnacles

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A community of grazing echinoderms have almost cleaned the bedrock surface of bryozoa and algae (depth 25 m). In the process they add fine bioclastic debris to the local sediment mass. Three Kings Islands, northernmost NZ. Image courtesy of Cam Nelson.

Echinoderms and barnacles are important contributors to bioclastic limestones

Echinoids are a diverse phylum with a geological history dating back to Early Cambrian. This entirely marine group of invertebrates includes the familiar Crinoids, star fish (Stelleroids and Ophiuroids), sea urchins (Echinoids) and sea cucumbers (Holothuroids), all of which have modern representatives. There are also several extinct groups. Most groups have five-fold skeletal division.

Devonian crinoid stems and disarticulated columnals in the bioclastic mix with fragmented fenestrate (f) and encrusting bryozoans (e).

Barnacles are arthropods, a phylum that includes spiders, scorpions, Horseshoe Crabs, and crustaceans like crabs, shrimp and lobster. They are included in this phylum because they have up to 6 appendages that, instead of being attached to a thorax and used as legs, are delicate, feathery wands that are extended and retracted in search of food. They are fully marine and as adults are firmly attached to rocky, or shelly substrates. The oldest definitive barnacles are Carboniferous, but it wasn’t until the Late Cretaceous that species diversity increased, becoming important contributors to Cenozoic cool-temperate limestones.

Although clearly disparate groups of invertebrates, echinoderms and barnacles do have a few attributes in common:
– Both are important contributors of bioclasts to cool-temperate carbonate sediments. They also play important roles in tropical reef systems, but the volume of sediment they produce is proportionally less.
– They shed calcium carbonate plates to the sediment mass; echinoderms also shed spines.
– They are bioeroders; barnacles erode or corrode the substrate (bedrock, other shelly organisms) prior to growing their basal plate, and echinoderms produce fine-grained bioclastic material while munching coral, bryozoa and calcareous algae.
– They are marine organisms that thrive in shallow environments washed by waves and currents, although some species occur at bathyal depths (e.g., some brittle stars; and barnacles like Bathylasma). They coexist with many other invertebrates and reef-builders.

Echinoderms

All common echinoderms (except the sea slugs) have segmented exoskeletons. Each segment, or plate is a single crystal of calcite; each attached spine is also a single calcite crystal. Calcite in modern echinoids is high in magnesium; the Mg is replaced by Ca during recrystallization to low magnesium calcite. When the animal dies, the exoskeleton breaks up into its constituent plates and spines.

Echinoderms bioclasts, even as small fragments, are relatively distinctive in thin section:

In life, the plates and spines are porous. In thin section this appears as uniformly distributed, microscopic perforations. The pores are filled with calcite cement during burial diagenesis; they may also be filled by glauconite.
In transverse sections, spines show a distinctive radial crosshatch pattern formed by the intersection of concentric growth rings and perforations arranged in radial fashion.
Plates and spines consist of a single calcite crystal that will show uniform extinction under crossed polars (unless the fragments have been strained during deformation).
Calcite cement overgrowths are usually syntaxial, such that both plate and spine plus cement will go into extinction simultaneously (they share the same crystallographic axes).
Crystallographic c axes parallel the long axes of spines.
The magnesium content of skeletal calcites is variable, even within a single specimen. For example, spines commonly vary from 5-10% MgCO₃, and plates up to 16%. Much of the Mg is replaced during diagenesis so that spines and plates plus their syntaxial cement overgrowths are low magnesium calcites (inclusions of HMC may be retained within the crystals).

Recent echinoderm, minus its spines that detached soon after the animal died. The spines were attached to the primary tubercles. The 5-fold symmetry is readily apparent. The five ambulacral and interambulacral areas each consist of interlocking calcite plates that tend to separate when the animal dies. Image credit Sheila Brown.

An abraded echinoderm plate in modern beachrock, Hawaii. The perforated structure is easily identified. Under crossed polars the grain extinguishes as a single crystal. Plain polarized light.

Recent echinoderm plate, slightly rounded by abrasion, sampled from temperate water bioclastic sand (that includes benthic forams (f), bryozoa, barnacle, and mollusc fragments). The perforations are unfilled. Left: Plain polarized light; Right: crossed polars with the grain in full extinction. Grain mount thin section. Three Kings Islands, northernmost NZ.

Grain mount thin sections of bioclastic sand sampled from about 30 m water depth, Three Kings Islands, northernmost NZ. Left: Small echinoderm spines with the basal tubercle attachment intact. Right: Longitudinal section of a broken and partly abraded echinoderm spine (s) showing the characteristic arrangement of perforations parallel to the spine long axis. Specimen occurs with unfilled foram tests (f) and abraded barnacle fragments (b). Both images plain polarized light.

Bioclastic beachrock (Hawaii) showing an abraded and bored echinoderm plate (top left). The margins of the plate and the borings (b) are lined with a thin rind of pale brown micrite. Two transverse sections of echinoderm spines (lower center and right) display typical radial fenestrate-like textures.

An Oligocene bioclastic, glauconitic limestone containing abundant abraded and broken echinoderm plates (e), most of which are encased in syntaxial calcite cement (s). Also present are numerous benthic forams (f), bryozoa fragments (b – some with glauconite filling), one small solitary coral (c – transverse section across septa), barnacle fragments with large pores filled by calcite cement (bc), and the occasional faecal pellet replaced by glauconite. Left: plain polarized light; Right: crossed polars. The syntaxial overgrowths extinguish at different stages for each echinoderm fragment.

Barnacles

Sketches of the two barnacle morphologies based on common modern forms. Left is the acorn barnacle species Balanus with operculum intact, attached by a basal calcite plate; Right is the common Goose barnacle Lepas that attaches to substrates with a soft pedicle that allows the animal to move with waves and currents.

All barnacles live attached to hard substrates (rock, shells, wood). The majority of species reside in shallow, agitated water although a few have evolved to survive the darkness at bathyal depths (e.g., Bathylasma). One group, the goose barnacles, are attached by a long leathery tube (a peduncle) to rocks, floating logs, and other flotsam. However, the most familiar species (the acorn barnacles) attach themselves with a basal calcite plate to rocks, other shells, whales, and ship’s hulls.

A typical acorn barnacle consists of robust calcite and aragonite wall plates fitted to the margins of the basal plate. Wall plates are commonly plicated, or folded, presenting a ribbed appearance. The front door, or operculum, consists of two smaller triangular shaped calcium carbonate plates (the scutum and tergum) that open during feeding.

When the animal dies, the operculum disaggregates. The wall plates may be overgrown by a new generation of barnacles or dislodged by waves or moving sediment to become part of the bioclast sediment mass.

A recent temperate to subtropical NZ gastropod, Cookia sulcata, covered by the common acorn barnacle, Balanus. Most of the barnacles have lost their opercula.

Live goose barnacles (Lepas) attached to a log that was washed ashore during a storm, their cirri searching vainly for food. The calcite plates will join the local sediment mass when the animals die or are predated by marauding sea gulls. Raglan, northwest New Zealand.

Apical view of a small, Pliocene acorn barnacle attached to a large oyster shell. The operculum and animal gut has been replaced by cemented carbonate sand. N.Z, Matemateaonga Formation.

Thin section cut transverse to a recent acorn barnacle plate, showing the array or pores at the base of the plate. Note the finely layered calcite that extends upwards into the characteristic plications, or folds (see image below). Plain polarized light.

A modern Balanid barnacle. Note the rib-like plications (left side, and the longitudinal pores exposed on the abraded plate (centre). Specimen from Parengarenga, northernmost New Zealand

Thin section cut approximately parallel to a recent acorn barnacle plate, showing the characteristic plicate (folded) calcite structure. The left margin corresponds to the base of the plate (cf. the image above). Plain polarized light.

Fragments of Oligocene barnacle plates (centre and top left) showing typical plicate structure. The central fragment has a thin, outer marginal layer; the upper fragment has remnants of the large basal pores. The plates are surrounded by small dark brown pelloids. Crossed polars. Image courtesy of Cam Nelson.

Acknowledgement

Many thanks to Kirsty Vincent, Earth Sciences, Waikato University for access to the petrographic microscope.

Carbonates in thin section: Molluscan bioclasts

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Molluscan bioclasts are important components of tropical, and cool-temperate water limestones

Bioclasts are probably the most important components of Phanerozoic limestone frameworks (grainstones, rudstones, packstones, floatstones). They include all manner of invertebrates that produce preservable shells and skeletons, algae that secrete calcium carbonate structures (particularly red algae), and sponges that maintain their rigidity by growing an interlocking scaffolding of calcite or silica. Some are microscopic (e.g., foraminifera, radiolaria, sponge spicules), others are macroscopic. They may be preserved intact or as broken, abraded fragments. The most common macroscopic invertebrate representatives include molluscs, corals, echinoids, cirripedes (e.g., barnacles), bryozoans, and calcareous algae.

Shell structure

By far the largest group of shelly invertebrates is the molluscs, represented by bivalves, gastropods, ammonoids and nautiloids, scaphopods, and amphineura (e.g., chitons). All secrete shells that are layered in calcite or aragonite; each layer is monomineralic but the mineralogy can vary among layers. Many species also grow a chitinous outer layer, the periostracum, but this has low preservation potential. Bivalves commonly have 2 to 4 layers, gastropods up to 6 layers. Each layer consists of microscopic crystals oriented normal, parallel, or oblique to the shell surface. The most common crystal fabrics are:

Prismatic, where high- or low-magnesium calcite or aragonite prisms are oriented normal to the shell surface. This is commonly the outer layer of a shell.

Prismatic aragonite forms the primary layer in this small gastropod. A very thin layer on the inside surface (arrow) could be either foliated or crossed lamellar aragonite. The shell whorl is filled by fibrous aragonite cement. Plain polarized light. Recent beachrock, Hawaii.

Foliated, where micron-sized calcite crystal plates create a fabric that approximately parallels the shell surface.

Part of a bivalve showing a thick layer of foliated calcite – basically a stack of calcite or aragonite plates. Plain polarized light.

Crossed lamellar structure where aragonite crystallites are organized in sheets. The crystallites in one sheet are oriented at relatively high angles to adjacent sheets. The overall effect can sometimes resemble a zig-zag pattern.

This rounded fragment in recent beachrock, could be either bivalve or gastropod. It is a good example of crossed lamellar structure (aragonite); the two sets of lamels are almost at right angles in this view. The main cement along the fragment margin is fibrous aragonite. Crossed nicols. New Caledonia.

Nacreous patterns of micron-sized aragonite plates arranged parallel to the surface – where this layer occurs on the inside of a shell, it is easily recognized by the lustrous, rainbow-like colours.

Most bivalves have two, three or four layers composed of calcite or aragonite, but there is much variation in structure depending on the species. For example, some species of the common mussel Mytilus contain a single, thick nacreous layer, whereas oysters (e.g., Ostrea) tend to consist of foliated calcite and a much thinner nacreous layer. Other species consist of combinations of layer types. Gastropods have up to 6 layers that in some species are all aragonite, and in others contain a thin layer of prismatic calcite. Scaphopods and cephalopods are composed entirely of aragonite.

Shells that contain aragonite are prone to recrystallization and replacement by sparry, low magnesium calcite. Mineral replacement usually destroys original crystal textures and may even obscure the distinction of individual layers, although traces of the former textures may persist. Examples of such textural changes are shown below.

Identification

Bioclasts are commonly fragmented. Fragment size is highly variable, depending on the original size of the critter, the degree of comminution resulting from bioerosion, and abrasion during transport and deposition, and to some extent diagenesis. In general, the larger the fragment, the easier the interpretation. However, there will be many fragments where even the phylum may be difficult to determine. Such problems are exacerbated by the (mostly) random orientation of a thin section slice through the bioclasts. The diagrams below show some of the characteristic morphological elements of mollusc shells that may help in this regard.

In thin section, the distinction among mollusc classes and species often relies on a few basic morphological elements that may be preserved in larger shell fragments.

Bivalves

Non-chambered. A change in shell curvature from ventral to dorsal margins is most pronounced at the umbo, although this may not be apparent in some species such as oysters. The thickened hinge area between two valves is located just below the umbo; if you are really lucky there may be evidence of dentition (teeth and sockets). However, small fragments can be difficult to distinguish from gastropod or even brachiopod fragments.

A small thin-shelled bivalve with hinged valves (hinge top left); valves thin towards the ventral margin. The outer layer (yellow-brown) has either crossed lamellar or foliated structure – this layer thins towards the umbo and is underlain by a (?) foliated layer. The inner surface is lined with a thin, isopachous aragonite rind. The sediment fill is predominantly pellets and sand. Plain polarized light. Holocene hardground, Abu Dhabi. Image courtesy of Stephen Lockier.

The ventral margin (i.e., opposite the hinged margin) of a bivalve in which the original carbonate has been recrystallized to calcite; note the neomorphic textures (n). The inner layer of (?) foliated carbonate is preserved in relicts outlined by microscopic inclusions of calcite, clay-sized sediment, and possibly altered organic matter. The shell paragenesis began with an initial rind of siderite cement (dark brown) followed by isopachous, scalenohedral calcite (arrow), and finally very coarse sparry calcite. Plain polarized light. Jurassic, northern British Columbia.

A Pliocene bivalve sectioned to show the hinge line (red arrow) and umbo. The original aragonite has been replaced by coarse neomorphic calcite spar, but remnants of the original foliated or crossed lamellar layers persist (black arrow). Plain polarized light. Image courtesy of Vincent Caron.

A partly recrystallized fragment of either a bivalve or gastropod showing relict foliated layering. This fragment is thoroughly bored, probably by sponges. The borings are filled by micritic calcite (dark brown). Top right is a bryozoa fragment. Plain polarized light.

Gastropods

Whorls are arranged in corkscrew fashion around a central, solid column (the columella); longitudinal and oblique sections will show this arrangement. Transverse sections will show the columella. The gastropod spire, whether short or long, will be tapered. Whorls will be filled with sediment and/or cements.

A small gastropod spire snuggled up against a much larger fragment of (probable) bivalve. The whorl cavities are lined initially by isopachous siderite (brown), followed by drusy calcite spar. The outer layer of the bivalve is mostly neomorphic calcite spar (n), but relict fabrics suggest an original crossed lamellar structure (arrow). The outer margin of the shell has numerous small borings (b). Plain polarized light. Jurassic, northern British Columbia.

A small, turreted gastropod. The main layer is prismatic aragonite. There is a thin foliated layer between each whorl. The whorl cavities are also lined with fibrous aragonite cement. In the bioclastic mix are echinoderm spines (e), barnacle fragments (b), and a couple of bivalve or gastropod fragments (f) with well-defined foliated structures. Plain polarized light. Recent beachrock, Hawaii.

A transverse section of a small gastropod (details of layering shown in image above). The central columella is characteristically thickened. The whorl cavity is filled with bundled isopachous fibrous aragonite cement. Its neighbours include a large benthic foram and an echinoderm fragment. Plain polarized light. Recent beachrock, Hawaii.

An oblique section through a gastropod, possibly a (pelagic) pteropod. The primary layer is cross lamellar structure. In the bioclastic mix are pellets plus small fragments of bryozoa and barnacles. Plain polarized light. Recent sediment, Three Kings Islands, northernmost New Zealand. Image courtesy of Cam Nelson.

Fragment of a large gastropod showing recrystallization of the original aragonite and calcite neomorphism to coarse spar. The top margin shows relict prismatic structure. The borings have been filled with host sediment. Plain polarized light. Provenance unknown.

Scaphopods

Also known as tusk shells, are tapered longitudinally and hollow down the centre where the animal resides. Transverse sections show a circular shell outline; the shell is layered as in other molluscs, but calcite replacement of the aragonite may obscure the original layer structures. The central part of the shell is hollow and will fill with sediment and or cements. In life, the shells are open at both ends.

Hand specimen view of a Jurassic seafloor littered with scaphopod shells plus a few small brachiopods and bivalves (Northern British Columbia). The shells show a crude current alignment. Cartoon on the right depicts a scaphopod in feeding position beneath the sediment-water interface. Apparently, they like to eat foraminifera.

An oblique section through a Jurassic scaphopod. Remnants of a crossed lamellar layer can be seen along the upper wall of the shell (arrow). The cavity has an initial thin rind of siderite (brown) cement (s), the remnants of an isopachous (i) cement (now calcite), and a later fill of neomorphosed calcite spar. Left, plain polarized light; Right, crossed polars. Northern British Columbia.

A transverse section of a scaphopod, replaced by neomorphosed calcite spar. The main layer shows remnants of crossed lamellar structure (arrow). The cavity has an isopachous veneer of acicular calcite (s), followed by a rim of small scalenohedral calcite crystals, with the remaining void filled by calcite spar cement. Part of the scaphopod fill presents a geopetal structure (g), where stratigraphic top was to the top right (yellow arrow); it consists of dark brown mud, quartz silt, and siderite. The fragment top right could be either mollusc or brachiopod – recrystallization of the calcite in this fragment has obliterated original defining textures. Left, plain polarized light; Right, crossed polars. Jurassic, northern British Columbia.

Distinguishing molluscs from other phyla

Brachiopods: Bivalved, shells always unequal but bilaterally symmetrical. Shells consist of calcite layers. Small fragments may be difficult to distinguish from molluscs, unless they contain punctae normal to the shell surface (punctae are tube-like perforations through the shell). Brachiopod shells are connected at a hinge but lack the dentition that is characteristic of bivalves.

Echinoderms: Echinoderm plates and spines consist of a single calcite crystal. Calcite cements are commonly syntaxial; this can be verified under crossed nicols where the plate (or spine) plus cement will move into extinction together. Plates and spines also contain micron-sized pores that will ultimately fill with syntaxial calcite cement, or minerals such as glauconite.

Barnacles: The interlaminate and plicate (fold) structures typical of barnacle plates (mostly low magnesium calcite) are diagnostic thin section features, even in small fragments. Individual plates may contain a mix of aragonite and calcite, unlike mollusc layers that are always monomineralic.

Foraminifera: Small gastropods sometimes resemble planispiral benthic forams. However, the chambers in forams are discrete and separate from their nearest neighbours; the body whorls in gastropods are continuous and arranged in corkscrew fashion around a central column.

Acknowledgement

Many thanks to Kirsty Vincent, Earth Sciences, Waikato University for access to the petrographic microscope.

Crystal axes, crystal systems, and crystal face intersections

Christian Weiss crystal face parameters

Intersection with one axis

Intersection with two axes

Intersection with three axes

William Miller intercedes

Miller-Bravais indices

Twin and cleavage planes

A couple of good resources

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Green schist

Biotite-hornblende gneiss

Biotite granite

Olivine basalt

Recrystallized chert

Silty mudstone

Some generalizations

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And in thin section…

Intercrystal boundaries in cements

Neomorphic fabrics

Aggrading neomorphism

Acknowledgement

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Common ignimbrite traits

Ignimbrites in outcrop

Ignimbrites in thin section

Acknowledgement

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Glauconite and glaucony

Glauconite (glaucony) in thin section

Chatham Rise

Modern sea floor, Three Kings Islands, northern New Zealand

Glauconite infill of fossils

Glauconitic micaceous arenite

Glauconitic arenite

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Foraminifera

There are four main orders

Sponges

Acknowledgement

Other links

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Differentiation from some other phyla

Acknowledgement

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Echinoderms

Barnacles

Acknowledgement

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Shell structure

Identification

Bivalves

Gastropods

Scaphopods

Distinguishing molluscs from other phyla

Acknowledgement

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