Optical thermal maturity parameters and organic

International Journal of Coal Geology 150–151 (2015) 74–119

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International Journal of Coal Geology journal homepage: www.elsevier.com/locate/ijcoalgeo

Optical thermal maturity parameters and organic geochemical alteration at low grade diagenesis to anchimetamorphism: A review Christoph Hartkopf-Fröder a,⁎, Peter Königshof b, Ralf Littke c, Jan Schwarzbauer c a b c

Geological Survey North Rhine-Westphalia, De-Greiff-Straße 195, 47803 Krefeld, Germany Senckenberg Research Institute and Natural History Museum, Senckenberganlage 25, 60325 Frankfurt am Main, Germany Energy and Mineral Resources Group (EMR), Institute of Geology and Geochemistry of Petroleum and Coal, RWTH Aachen University, Lochnerstraße 4–20, 52056 Aachen, Germany

a r t i c l e

i n f o

Article history: Received 6 November 2014 Received in revised form 14 June 2015 Accepted 15 June 2015 Available online 19 June 2015 Keywords: Vitrinite Conodonts Palynomorphs Invertebrates Thermal maturity Organic geochemical alteration

a b s t r a c t Sedimentary organic matter derives from biological precursor material which undergoes systematic, irreversible changes upon burial which mainly reflect increasing diagenetic temperatures, although other factors have also an influence. Several parameters have been established in the last decades to determine the palaeotemperature history of sedimentary rocks based on geochemical or petrographical methods. Organic matter is the most temperature-sensitive solid constituent in sedimentary rocks and vitrinite reflectance (VR), miospore and conodont colour alteration are among the most widely used optical maturity parameters. Despite tremendous interest in estimating maturity parameters in the oil generation zone as well as for high grade diagenesis and metamorphism and despite decades of research only a few of these methods have been well established and compared to each other by now. The focus of this review is on some new aspects with respect to organic matter maturation and optical palaeotemperature parameters, especially for high grade diagenesis and anchimetamorphism. Such a discussion might be a prerequisite for a better understanding of palaeotemperature assessments in different sedimentological and/or geotectonic settings and useful for different fields in applied sciences. Furthermore, suggestions for further research are discussed. © 2015 Elsevier B.V. All rights reserved.

1. Introduction During the last decades, reconstruction of palaeotemperature histories of sedimentary rocks has become an important discipline in earth sciences, which is widely applied in petroleum exploration. Whereas the present-day temperature regime in sedimentary basins can be studied on the basis of borehole temperatures or temperature logs, the reconstruction of palaeotemperature evolution is more difficult. In many sedimentary basins, present temperatures are much lower than ancient temperatures and the state of the rocks with respect to e.g. petroleum generation and compaction was established during earlier geologic eras at higher temperatures. Especially for these basins, understanding of the palaeotemperature histories is critical and a prerequisite when quantifying diagenesis or modelling petroleum generation. However, palaeotemperature reconstructions are essential not only for basins and sedimentary rocks which experienced their highest temperatures in the past, but also for those in which Neogene temperatures are the highest. This is because early temperature evolution will have ⁎ Corresponding author. E-mail addresses: [email protected] (C. Hartkopf-Fröder), [email protected] (P. Königshof), [email protected] (R. Littke), [email protected] (J. Schwarzbauer).

http://dx.doi.org/10.1016/j.coal.2015.06.005 0166-5162/© 2015 Elsevier B.V. All rights reserved.

already influenced mineral precipitation and oil or gas generation from source rocks. The knowledge of the extent of such an early phase of generation can be a clue towards an understanding of the extent of the late (Neogene) phase of generation in different parts of a basin and thus an aid in petroleum exploration strategies. The irreversible changes of organic matter reflect to a great extent the increasing diagenetic temperatures (e.g. Price, 1983), although other factors such as microbial degradation (at shallow depth) and pressure have also an influence. The temperature-related changes can be measured using a variety of optical and chemical maturity parameters. Maturation is a term commonly used in sedimentary basin studies to address thermally induced changes in the nature of organic matter. Maturation reflects organic matter conversion including petroleum generation and thermal gas generation. The driving force is the difference in free energy (− ΔG) between the reactant (immature organic matter) and the product (mature organic matter; see Atkins and de Paula, 2010). Therefore, maturation depends mainly on temperature and the time, during which specific temperatures are maintained. Other factors such as the chemical environment and pressure are generally regarded to be of lesser importance, although there may be exceptions (Carr, 1999; Huang, 1996; Price and Barker, 1985). The correlation of optical measurements and maturity is based on the thermodynamically driven changes of the molecular composition.

C. Hartkopf-Fröder et al. / International Journal of Coal Geology 150–151 (2015) 74–119

This linkage is obvious especially for the systematic changes of the organic matter. Optical properties in the range of UV and visible light are related to the bond configurations and the corresponding quantized energy levels of the valence electrons. Highly unsaturated and conjugated systems and in particular aromatic moieties, both with delocalized πelectron systems, build up chromophores in the visible light range. With increasing temperature a higher tendency to form aromatic systems of higher cyclization degree can be observed in natural organic matter. A higher aromatic cyclization degree results in a lower energy range of absorption. Hence, with increasing temperature and corresponding increasing aromaticity a shift in visible light absorption from the high energy range to the lower energy range has to be expected. For predicting the colour changes with temperature two further factors have to be considered: the complementary colour system affecting human colour perception and the superimposition of different colours resulting from lower and higher energy absorption at a later stage of maturity. Regarding all these aspects, a shift from light yellow colour towards orange tones at initial state followed by red/brown to dark colour tones resulting from the ongoing superimposition of colours can be expected. In the same way, also fluorescence properties are affected systematically by the changes of the molecular composition with increasing temperature. In order to quantify thermal maturation of sedimentary rocks, a great number of optical, other physical and chemical maturity parameters have been developed. All these parameters are measured either on total organic matter or parts of the organic matter. Table 1 gives an overview on some of the most common parameters used and their application. The most widely used parameters in the petroleum industry are vitrinite reflectance, miospore colour, and Tmax values. Furthermore, molecular geochemical parameters are applied, which are especially useful in oil–oil and oil–source rock correlations. Many of the maturity parameters listed in Table 1 (based on a compilation by Littke et al., 2008a) have been developed for the oil generation zone. Usually, the main zone of oil generation is defined to range from about 0.5–0.6% VRr (random vitrinite reflectance; oil birth line; Rullkötter et al., 1988) up to 1.3% VRr (oil death line). It should be noted, however, that petroleum generation depends as well on kerogen type. For example, sulphur-rich kerogen can generate (heavy) bitumen already at low temperature and maturity (Orr, 1986). In terms of maturation stages kerogen in the main zone of oil generation is defined as mature (oil window), below 0.5–0.6% VRr as immature, between 1.3% VRr and 2.0% VRr as postmature and above 2.0% VRr as overmature (Table 1). In coal petrography, often the term “coalification” is used describing the evolution of peat towards lignite, subbituminous coal, bituminous coal, anthracite and finally meta-anthracite. Based on organic maturation, an ever increasing number of parameters have been developed over the last decades, providing insights into temperature histories or maximum palaeotemperatures that sedimentary rocks experienced (Harris and Peters, 2012; Peters et al., 2005; Suárez-Ruiz et al., 2012; Taylor et al., 1998). Among optical maturity parameters, vitrinite and huminite reflectance are most commonly used. However, optical and chemical properties of other macerals, palynomorphs, and microfossils with organic compounds embedded in their shell/skeleton, and molecular composition of bitumen also serve as parameters of maturity (Table 1). More parameters have been proposed, e.g. colour changes in foraminifera tests (Foraminifera Colouration Index, FCI, e.g. McNeil et al., 1996), ostracod shells (e.g. Kontrovitz et al., 1992), conchostracan valves (Tasch, 1982), other arthropod cuticles (Bartram et al., 1987), and in ichthyoliths (e.g. Johns et al., 2012). In addition, colour changes induced by the effect of heat have been reported from other fossil groups e.g. thecamoebians (McNeil et al., 2000), microbivalves, microgastropods (Ainsworth et al., 1990), and amniote eggshells (Janssen et al., 2011). However, due to the scarcity of these fossils or because their potential as palaeogeothermometer has not yet been elucidated in detail they have not found application in thermal maturity studies. Furthermore, colour

75

alteration of calcareous shells must not inevitably be coupled with high temperatures but may be a taphonomic indicator for longer exposure times before burial (Kolbe et al., 2011). For the oil window, the change in optical properties of vitrinite/huminite, solid bitumen, conodonts, graptolites, acritarchs and algae/alginite, miospores (spores and pollen grains)/sporinite, chitinozoans and scolecodonts has been investigated partly in great detail (e.g. Bertrand, 1990a; Bertrand and Malo, 2012; Epstein et al., 1977; Goodarzi and Norford, 1985; Hoffknecht, 1991; Petersen et al., 2013; Schoenherr et al., 2007; Staplin, 1969; Taylor et al., 1998). Many cross-correlation charts have been published showing various organic maturation indicators such as vitrinite reflectance (VR), conodont colour alteration (CAI), miospore colour (TAI, SCS, SCI), miospore fluorescence, often in correlation with hydrocarbon generation (e.g. Batten, 1996b; Harris et al., 1987; Kovács and Árkai, 1987; Traverse, 2008; Taylor et al., 1998). These charts offer a rough correlation of thermal maturity parameters, but are usually limited to few indices. A perfect correlation cannot be reached, because the progress of the various chemical and physical changes follows different kinetics. Generally correlation between different methods is difficult, particularly in defining the transition between diagenetic to anchizone. After Kovács and Árkai (1987) the transition between diagenesis and anchizone corresponds to CAI 5. A comparative study of CAI and illite crystallinity (IC) by García-López et al. (1997) proposes that the transition between diagenesis and anchizone (180–230 °C, e.g. Kisch, 1990; Frey and Robinson, 1999) corresponds to CAI 4, and the transition between anchi- and epizone (280–320 °C after Bucher and Frey, 1994) occurs at about CAI 5.5. Kovács and Árkai (1987) have correlated CAI and IC in a petrologically heterogeneous set of metamorphic rocks from Hungary. They correlate the transition from diagenesis to anchizone with CAI 5, i.e., with a higher CAI than by García-López et al. (1997). The purpose of this paper is to review the extensive data on the most useful optical palaeotemperature parameters determined taking particularly also high grade diagenesis to anchimetamorphism into account which has been investigated much less than low grade diagenesis/oil window stage. Also chemical composition and structure of organic matter are treated to some extent, especially in relation to changes in optical properties. Although there exists a huge number of publications dealing with different aspects mentioned above, several open questions still remain which will be addressed. 2. Palaeotemperature parameters 2.1. Vitrinite and solid bitumen reflectance in comparison to geochemical maturity parameters and fission tracks Arguably the most accurate way to reconstruct temperature histories in the context of sedimentary basin history is numerical modelling (basin and petroleum system modelling; e.g. Littke et al., 2008b). Clearly, the quality of model predictions on palaeotemperature, maturation, petroleum generation etc. depends on the availability of temperaturesensitive data and parameters, which can be used for calibrating these models. The most widely used of these parameters is vitrinite reflectance. Vitrinite is a group of organic particles derived from higher land plants (“wood-like particles”, “huminite” in lignites) which are the major constituents of most coals, but also ubiquitous in other sedimentary rocks. Its chemical properties as well as its reflectance change systematically with increasing burial temperatures. Vitrinite reflectance was originally measured on coals, where vitrinite is the most abundant, usually predominant constituent. Later, vitrinite was found to be common in other sedimentary rocks as well, especially in dark-coloured siltstones, sandstones, marlstones and shales. This abundance allowed a wider application as maturity parameter predicting e.g. the stage of oil generation for petroleum source rocks. In particular, it served as calibration tool in numerical modelling of thermal histories of sedimentary basins. The basics of vitrinite reflectance are described in detail in

76


Table 1 Listing of various thermal maturity parameters. * Onset of petroleum generation depends on kerogen type. ** Should not be mistaken for TOC, Corg.

Organic geochemistry

Colour

Fluorescence

Reflectance

Maturation parameter

Short term(s)

Method

Application

Hydrocarbon maturity stages

Vitrinite reflectance

Optical, selected woody organic particles

VR r; R r; R o

Devonian and younger; all sedimentary rocks

Immature (0.2 – 0.6 %)*, mature (0.6 – 1.3 %), postmature (1.3 – 2.0 %), overmature (> 2.0 %)

Maximum vitrinite reflectance

Optical, selected woody organic particles

VR max ; R max

Devonian and younger; all sedimentary rocks

Semi-anthracite (2.2%) to metamorphism (> 6%)

Solid bitumen reflectance

Optical, selected organic particles

BR r

All sedimentary rocks, preferentially source, carrier and reservoir rocks

Mature to overmature

Graptolite reflectance and bireflectance

Optical, graptolites

GR; R grap

Late Cambrian to Carboniferous; marine sedimentary rocks

Immature to overmature

Chitinozoan reflectance

Optical, chitinozoans

CR; R chi

Cambrian to Devonian; marine sedimentary rocks


Scolecodont reflectance

Optical, scolecodonts

SR; R sco

Cambrian and younger; marine sedimentary rocks


Hydroid reflectance

Optical, hydroids

HR; R hyd

Ordovician and younger


Visible fluorescence colour and intensity

Optical, selected organic particles, e. g. miospores, algae (dinoflagellates, prasinophyceans), acritarchs

λmax, Q , FIint

All sedimentary rocks, preferentially source rocks; also oil inclusions

Immature to postmature

Conodont colour

Optical, conodonts

CAI

Cambrian toTriassic; T marine sedimentary rocks; preferentially limestones or marbles and unconsolidated marls


Foraminifera colour and mineralogy

Optical, foraminifera

FCI, zone A – D

Cambrian and younger; marine sedimentary rocks


Late Ordovician and younger

Immature to postmature

Devonian and younger

Immature to ?

IAI


Immature to ?metamorphism

Optical, amorphous organic matter

TCI

Precambrian and younger


Acritarch colour

Optical, acritarchs

AAI, RGB colour space

Mesoarchean and younger

Immature to marginally overmature

Dinoflagellate colour

Optical, dinoflagellates

Mid–Late Triassic and younger

Prasinophycean algae colour

Optical, prasinophycean algae

Follows e. g. TAI, SCI AAI

Precambrian and younger

Immature to ?marginally overmature Immature to marginally overmature

Miospore colour

Optical, miospores, cryptospores

e.g.TAI, SCS, SCI, stTAI, PDI, % St , CIE and RGB colour space

Ordovician and younger; preferentially non-marine and marginally marine sedimentary rocks


Arthropod cuticle colour

Optical, arthropod cuticles


Immature to ?

Carbon Preference Index

Geochemical, specific non-aromatic hydrocarbons

CPI

All sedimentary rocks, unweathered material required

Immature to mature

Methyl Phenantrene Index

Geochemical, specific aromatic hydrocarbons

MPI


Early mature to slightly overmature

„Biomarkers” (various parameters)

Geochemical, specific non-aromatic or aromatic hydrocarbons

See Peters et al. (2005) and Walters et al. (2012)


Immature to mature

Rock-Eval T max

Geochemical, total organic matter

T max

All sedimentary rocks

Immature to slightly overmature

Rock-Eval PI


PI

A ll s edimenta ry ro cks , unweathered material required

Immature to slightly overmature

Rock-Eval HI


HI

A ll s edimenta ry ro cks

Immature to slightly overmature, depends on petroleum generation potential and maturity

Volatile Matter Yield

Geochemical, total organic matter (water- and ash-free)

VM

Only coals


Water content/moisture

Geochemical, total organic matter (ash-free)

Only coals

Immature to early mature

Carbon content**

Geochemical, total organic matter (water- and ash-free)

Only coals


Ostracod colour

Optical, ostracods

Conchostracan colour and mineralogy

Optical, conchostracans

Ichthyolith colour

Optical, ichthyoliths

Amorphous organic matter colour

– –

–

– C**


several textbooks (e.g. Taylor et al., 1998) as well as review articles (Suárez-Ruiz et al., 2012) and will not be repeated here. From a chemical point of view, increase of vitrinite reflectance seems to be related to an increase in aromaticity and condensation of aromatic moieties (Amijaya and Littke, 2006; Carr and Williamson, 1990; Dyrkacz et al., 1984; Ibarra et al., 1996; Kuehn et al., 1982; Schenk et al., 1990; Wollenweber et al., 2006) which has also been proven by μ-FTIR spectroscopy studies (al Sandouk-Lincke et al., 2013; Mastalerz and Bustin, 1993). Fig. 1 shows the ratio of aromatic stretching band areas and vitrinite reflectance as a function of pyrolytic temperature for pure vitrinite (modified from Schenk et al., 1990). Clearly, increase of reflectance correlates well with increase of condensed aromatic ring systems. Petersen and Vosgerau (1999) demonstrated a well-defined suppression of vitrinite reflectance with increasing resinite content in coals from eastern Greenland. Likewise, Littke (1987) pointed out differences in vitrinite reflectance for different types of vitrinite within single coal seams; in particular, vitrinites associated with cutinite and resinite showed low reflectance. Furthermore, very sulphide-rich coals had lower vitrinite reflectance as compared to “normal” coals indicating an influence of early diagenetic sulphate reduction on vitrinite structure. Chen et al. (2013) used μ-FTIR spectroscopy mapping of different macerals. Interestingly, vitrinite adjacent to resinite displays a more aliphatic structure as compared to more distant vitrinite, “suggesting that chemical

1.5

a) VR r [%]

1.0

0.5

0.0 0

100

200

300

400

500

600

Temperature [°C] 15

A 1600/A 1500

b) 10

components from resinite can diffuse over short distances into adjacent vitrinite, specifically causing hydrogen enrichment” (Chen et al., 2013). The area influenced by the adjacent maceral seems to be very narrow. Nevertheless, intermaceral effects seem to occur during the peat forming stage or possibly even later during thermal maturation and petroleum generation. Also, differences between vitrinite reflectance values measured on coals as compared to other sedimentary rocks have been reported (e.g. Bostick and Foster, 1975). Goodarzi et al. (1988) found for low maturity levels slightly higher vitrinite reflectance values for carbonates and lower ones for shales as compared to coals. Scheidt and Littke (1989) studied a large number of bituminous coals and adjacent mudstones, siltstones and sandstones. They found the same general trend, but much more scatter for the clastic rocks. Values were on average slightly higher in the coals (by 0.02%) for a vitrinite reflectance range from 0.6 to 1.5%. Whereas some studies revealed even slightly higher reflectance values in clastic rocks than in adjacent coals (Jasper et al., 2009), others reported a stronger suppression of vitrinite, especially in liptinite-rich petroleum source rocks (Hutton and Cook, 1980; Kalkreuth and Macauley, 1984; Petersen et al., 2006; Wenger and Baker, 1987), or retardation of reflectance increase under pore overpressure conditions (Carr, 1999). Barker et al. (2007) compared vitrinite reflectance of extracted and not extracted samples, showing that no difference exists. Furthermore, all bituminous coals are rich in solvent extractable organic matter including abundant hydrocarbons (Littke and Leythaeuser, 1993; Littke et al., 1989; Wilkins and George, 2002) without showing any evidence of suppression of reflectance. Thus the presence of bitumen/hydrocarbons does not seem to have a significant effect on vitrinite reflectance. In contrast, primary composition, e.g. wood rich in plant waxes leading to perhydrous vitrinite, and early diagenetic processes might create differences in reflectance. Large differences can occur between different lithologies at high rank due to the development of vitrinite anisotropy. Especially cleavage domains in shales can show vitrinite with high anisotropy and high mean and maximum reflectance as compared to adjacent microlithons (Littke et al., 2012). Therefore at mean vitrinite reflectance levels greater than 2–3%, lithology description especially with respect to cleavage becomes indispensable. A literature overview on lithology effects on vitrinite reflectance has recently been published by Suárez-Ruiz et al. (2012). Thus, the evolution of vitrinite reflectance as a function of temperature and time has been intensely studied for more than 50 years and used in sedimentary basin numerical models as calibration tool for about 40 years (Lopatin, 1971; Waples, 1980). At present, calculation of temperature histories from vitrinite reflectance data is mainly based on the algorithm published by Sweeney and Burnham (1990): VRr ¼ eð‐1:6

5

0 0

100

200

300

400

500

600

Temperature [°C]

Fig. 1. Temperature dependence of a) vitrinite reflectance at heating rates of 0.1 (blue line) and 2.0 °C/min (red line) and b) the ratio of A1600/A1500 aromatic ring stretching band areas. The values of A1600/A1500 for anthracene, pyrene and coronene are shown for comparison; modified from Schenk et al. (1990). Note that results of well-defined laboratory experiments on a xylite sample are displayed here in order to demonstrate the link between reflectance and chemical structure of vitrinite; under natural conditions, development of vitrinite reflectance and aromaticity increase occurs at much lower temperatures and heating rates.

77

þ 3:7 FÞ

ð1Þ

In this equation, F is a stochiometric factor ranging from 0 to 0.85 and VRr is the mean vitrinite reflectance measured in oil immersion on randomly orientated grains. Using the method of Sweeney and Burnham (1990), vitrinite reflectance can be calculated, if temperature history is known. In many sedimentary basins, maximum temperatures were reached in the past – in these cases, the burial and heat flow conditions during times of maximum temperature can be deduced from the comparison of measured and calculated vitrinite reflectance. Fig. 2 shows an example of this procedure for a case, where numerous vitrinite reflectance data are known from various depths within a well. A high maturation/coalification at the surface neither proved deep burial followed by erosion nor the presence of high heat flow, e.g. due to magmatic intrusions. However, a strong vitrinite reflectance gradient indicates high heat flows (Fig. 2c, d), whereas a weak gradient clearly points towards deep burial (at moderate to low heat flows; Fig. 2b) followed by erosion.

78


a) time gap

b) best fit model Time [Ma]

Carbon.

Permian Triassic Jurassic Cretaceous ”Tertiary”

300

200

100

Easy%R o 0

0

2

4

6

0[m]

Burial depth

0[m]

1000

2000 5000 3000

4000 10 000

5000

Calibration

150 100 50

own measurements external measurements

Heatflow (mW/m 2 )

c) high heat flow model Time [Ma] 300

200

100

d) igneous intrusion Easy%R o

0

0

2

4

Time [Ma] 6

200

100

Easy%R o 0

0

2

4

6

Burial depth

0[m]

Burial depth

0[m]

300

5000

5000

10 000

10 000

150 100 50

150 100 50

Heatflow (mW/m 2 )

Heatflow (mW/m 2 )

Fig. 2. Principle of numerical basin modelling. Due to the time gap between the Jurassic and Paleogene (a) a model of sedimentation and erosion had to be developed calibrated by vitrinite reflectance data (b) indicating that deep burial (and high temperatures) occurred during the Cretaceous. The assumption of a high (c) or even very high heat flow (as typical of igneous intrusion; d) and lower burial depth leads to an overestimation of thermal maturity (c, d); modified from Senglaub et al. (2006). Therefore vitrinite reflectance vs. depth trends are a key to understand both heat flow and thickness of eroded sediments, especially in combination with fission track data revealing information on timing of uplift.

The burial and temperature modelling described above is generally based on calibration using the Sweeney and Burnham (1990) algorithm. However, some researchers argue that under “normal” burial conditions there is always sufficient time available for vitrinite particles to adapt and that, therefore, vitrinite reflectance can be directly translated into maximum palaeotemperature (Barker and Pawlewicz, 1994). They developed Tpeak ¼ ðlnVRr þ 1:68Þ=0:0124

ð2Þ

for “normal” burial conditions (slow heating) and for hydrothermal

conditions (rapid heating): Tpeak ¼ ðlnVRr þ 1:19Þ=0:00782

ð3Þ

One strength of this approach is the simple calculation of maximum palaeotemperatures, one weakness is that just two heating rates are considered rather than a wide spectrum of heating rates existing in nature. Not only organic matter, but also clay minerals react in response to temperature and pressure conditions within sedimentary basins. In particular the transformation of smectite into illite and the increase of


IC have long been used to describe these transformations (e.g. Frey et al., 1980; Lanson and Besson, 1992). Although these transformations depend on chemical environment, i.e. potassium availability, correlations with vitrinite reflectance have been published. A recent summary by Ferreiro Mählmann et al. (2012) indicates that at high heat flows (geothermal gradients) organic matter transformation is much faster than clay mineral reaction, i.e. vitrinite reflectance values are high but IC can remain low. In contrast low geothermal gradients have the opposite effect (Fig. 3). Thus clay mineral transformations can add very valuable information on thermal histories of sedimentary rocks and basins, especially at high grade diagenesis or low grade metamorphism, but should ideally be accompanied by organic matter maturation data. For rocks which reached maximum burial in the past and were later affected by significant cooling, arguably the best temperature history reconstructions are gained by a combination of vitrinite reflectance and fission track data as well as numerical basin modelling. Fission track thermochronology has been developed as a radiometric method based on the spontaneous fission of 238U (e.g. Wagner and van den Haute, 1992). Two fission fragments move in opposite directions through the crystal lattice causing linear zones of lattice damage. These fission tracks are affected by rock temperatures, leading to complete annealing at high, partial annealing (shortening) at intermediate, and preservation at

Integrated data area from different Helvetic and Austroalpine nappes

0.6

a)

Samples with regular vitrinite bi-reflectance Samples with pyro-vitrinite and bituminite close to intrusions, Vichuquén (Belmar et al. 2002)

0.5

Diagenesis Se -

rin ve

0.4

600 °C e 20 pp al. na et ea u st lav su iu Co (C

Anchizone 0.3

0.2

Epizone

)

1.0

2.0

Danubian s va nappes et lley South Carpathians al (F .1 r (Ciulavu 98 ey 0) et al. 2008)

Samples from the Monte Mattoni – Blumone (Gabbro) Samples from the Val Fredda (Quartzdiorite)

3.0

4.0

5.0

6.0

7.0

8.0

9.0

10

Vitrinite reflectance VRmax [%] 0.7

Orogenic Diagenesis and Metamorphism

b)

HyperLow thermal geothermal gradient gradient

0.6

ð4Þ

800 °C

us

08

0.0

low temperatures (Fig. 4). Tracks are usually studied in apatite, zircon, and titanite. The calculation of an age, representing the time of cooling below a specific threshold temperature, is primarily based on the ratio of U content and number of tracks accumulated within a certain volume of a crystal. It should be noted that there is no sharp closing temperature for fission tracks, but instead a temperature range of partial track stability. Within this range (called PAZ, partial annealing zone, Fig. 4) the length of the tracks is successively shortened. Thus, track length distributions represent specific temperature histories. One important aspect of apatite fission tracks is that the PAZ roughly represents the oil window; therefore temperature histories relevant for oil generation can be reconstructed and even influence exploration strategies. In contrast, zircon fission tracks are relevant for high grade diagenesis and low grade metamorphism (Fig. 4). Vitrinite (huminite) reflectance ranges from 0.2 to 0.6% in the immature stage which is roughly equivalent to the peat, lignite, and subbituminous coal stage, from 0.6 to 1.3% in the mature petroleum generation stage, from 1.3 to 2.0% in the wet gas stage and is above 2% in the dry gas stage equivalent to anthracite rank. Aside from vitrinite reflectance, many other optical maturity parameters have been developed (Table 1). Solid bitumen reflectance has often been applied, since this type of particle is common in many rocks in petroleum-bearing basins. Problems with respect to the application of this parameter have recently been discussed by Petersen et al. (2013). Solid bitumen is derived from decomposition (cracking) of former oil or is an early generation product of kerogen decomposition and is accordingly not only most common in petroleum source and reservoir rocks, but also in former carrier rocks. These rocks are often poor in vitrinite; therefore reflectance of solid bitumen can act as an alternative maturity parameter there. It also occurs in rocks older than Devonian in which true vitrinite particles are absent due to the lack of higher land plants. Correlation between vitrinite and solid bitumen reflectance is visualized in Fig. 5a (Schoenherr et al., 2007) and can be expressed as VRr ¼ ðBRr þ 0:2443Þ=1:0495

Re

0.1 Measurement limit of the device control

79

In rocks older than Devonian, vitrinite-like particles are commonly observed. Based on a comparison of reflectance values with those of different zooclasts and solid bitumen, Petersen et al. (2013) concluded that these vitrinite-like particles might be derived to a great extent from fragments of graptolites. Further optical maturity parameters are listed in Table 1. In particular, miospore and conodont colour as well as fluorescence parameters can be regarded as rapid indicators for thermal maturation. However, colour/fluorescence depend also on other factors rather than just temperature, e.g. on the thickness of the fossil or on the degree of oil impregnation (see below). New approaches

0.5

Diagenesis

Moderate (normal) geothermal gradient

0.2

Epizone High geothermal gradient

0.1 Measurement limit of the device control

0.0

1.0

2.0

3.0

4.0

5.0

6.0

7.0

180°C

PAZ zircon

increase of strain

380°C 8.0

9.0

event ng at i e H

ow Sl ) b

10

o co

g lin

60°C

Apatite 120°C

Number of tracks

0.3

c) Anchizone

a) Fast cooli ng

Hypothermal gradient

c)

Shift of the gradient due to contact metamorphism

Temperature

0.4

b)

a)

Vitrinite reflectance VR max [%]

Time Fig. 3. Kübler-Index/vitrinite reflectance plots. (a) Comparison of Kübler-Index/vitrinite reflectance correlation studies compiled for orogenic settings with hyper-thermal extensional and contact metamorphism; modified from Ferreiro Mählmann et al. (2012). (b) Conclusive Kübler-Index/vitrinite reflectance plot showing strong geothermal (heat flow) dependent trend-evolutions regarding the geodynamic setting; modified from Ferreiro Mählmann et al. (2012).

Track length

Fig. 4. Schematic illustration of length distributions of fission tracks resulting from thermal histories passing in specific ways the partial annealing zone (PAZ); modified from Littke et al. (2008a).

80


0.0

0.0

a)

b)

c)

0.5

0.5

Oil Window

Oil Window

Oil Window

1.0 1.5

1.5

Type I (G.R.S.)

VR r [ % ]

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2.0

y = 0.60x + 0.37 r = 0.96 n = 16

Type III (various coals)

3.0 400

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500

550

T max [ °C ]

3.5

VR r [ % ]

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2.5 3.0

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1.7

Methylphenantrene Index ( MPI )

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y = 1.0495x - 0.2443 4.5 5.0 0.0

G.R.S.

R2 = 0.9403

Green River Shale Data used for equation

0.5

1.0

1.5

2.0

2.5

3.0

3.5

4.0

4.5

5.0

BRr [ % ]

Fig. 5. (a) Correlation of vitrinite reflectance with solid bitumen reflectance (Schoenherr et al., 2007). (b) Tmax values from Rock-Eval pyrolysis (Taylor et al., 1998; note that trend for type II kerogen is similar to the one for type III). (c) MPI for type III kerogen (modified from Littke et al., 2008a).

have recently been developed to study maturation based on optical properties of vitrinites and other macerals, i.e. fluorescence alteration of multiple macerals (FAMM; Kalkreuth et al., 2004; Wilkins et al., 1995), μ-Raman spectroscopy (e.g. Wilkins et al., 2014) and μ-FTIR spectroscopy approaches (Chen et al., 2013). Excellent maturity parameters exist for coals, including vitrinite reflectance, water content, volatile matter yield, carbon content, and calorific value (Table 1). With the exception of vitrinite reflectance, none of the above parameters can be utilized on rocks other than coal. In contrast, there are several geochemical parameters applicable for different rock types. In particular, Rock-Eval parameters Tmax (temperature of maximum pyrolysis yield) and PI (Production Index) have often been used in the petroleum industry. PI values depend on the progress of petroleum generation and tend to increase with depth/temperature, but are also influenced by petroleum impregnation or expulsion. The more widely used Tmax values increase with depth/temperature. They have been correlated with vitrinite reflectance for terrestrial organic matter (Type III kerogen) and humic coals (Petersen, 2006; Petersen et al., 2013; Teichmüller and Durand, 1983; see Fig. 5b), but their development with increasing temperature depends greatly on thermal stability of organic matter, i.e. kerogen type. Nevertheless, Tmax values in combination with Hydrogen Index (HI) and PI values from Rock-Eval pyrolysis have excellent potential to predict quality of organic matter (kerogen) with respect to petroleum generation potential and maturity. Pitfalls of the technique have been discussed in detail, e.g. by Dahl et al. (2004) and Peters (1986). In particular, the presence or absence of specific minerals can change Rock-Eval parameters (e.g. Jasper et al., 2009). Comparison of Rock-Eval parameters with a variety of other coal rank parameters has been published by Bostick and Daws (1994). In addition, there are many molecular geochemical parameters which are well suited to study thermal maturity, especially for the range of maximum temperatures between 50 and 150 °C. These parameters are usually less suitable for weathered rocks from outcrops and for high ranges of maturity, exceeding 1.3% VRr. Exceptions are diamondoid-based

parameters (Böcker et al., 2013; Chen et al., 1996; see Fig. 6) and the Methylphenantrene Index (MPI; Fig. 5c) established for terrigenous Type III kerogen by Radke and Welte (1983), which cover a greater maturity range than other molecular parameters. Also widely applied is the “Carbon Preference Index” (CPI) value which reflects the ratio of oddnumbered over even-numbered n-alkanes. In addition, there are many parameters which depend on ratios of similar molecules, in particular stereo- and structural isomers, which have, however, different thermal stabilities (biomarkers). Details on geochemical maturity parameters can be found in e.g. Peters et al. (2005) and Walters et al. (2012). Clearly, the best way to decipher maturation is a combination of different parameters, and in most cases a combination of optical and organic geochemical indices. 2.2. Conodonts 2.2.1. General remarks Conodont elements are tiny (0.2–5 mm) remains of feeding apparatuses of an extinct group of probably primitive vertebrates (e.g. Aldridge and Donoghue, 1998; Briggs et al., 1983; Knell, 2013; Lindström, 1964; Sweet, 1988; Szaniawski, 2009). Based on a recent paper by Turner et al. (2010) this assumption seems questionable but conodonts have to be considered as basal chordates (Blieck et al., 2009). Independently of their palaeontological affinity conodont remnants are very common in marine Cambrian to Upper Triassic sedimentary rocks. Conodonts occur in a big variety of shapes and owing to their rapid evolution, widespread distribution, and relative abundance they are used as excellent stratigraphic tools providing a high-resolution biostratigraphic framework for the Palaeozoic and Triassic time (Sweet, 1988). Although found in a range of sedimentary and even metamorphic rocks such as marbles, conodont elements are most easily recovered from carbonate rocks. In most sequences, conodont abundance is inversely proportional to the rate of deposition of their host rock. Conodont elements are composed of a calcium phosphate mineral, namely carbonate fluorapatite (Olcott Marshall et al., 2013; Pietzner


DM AI 0

20

40

60

M AI 80

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30

40

50

60

70

81

M DI 80

90 100

20

30

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50

60

70

80

0.5

Lithology Squares: Coal

1.0

Circles:

Organic-rich sedimentary rock

VR r [%]

1.5

Coal seam names (Ruhr Basin) V K P H M D

2.0

2.5

Ibbenbüren South Wales 3.0

Upper Silesian Coal Basin

3.5

Suggested trendline 0

Fig. 6. Correlation of vitrinite reflectance with diamondoid maturity parameters DMAI, MAI, MDI; modified from Böcker et al. (2013); see Chen et al. (1996).

et al., 1968; Trotter and Eggins, 2006; Trotter et al., 2007). Conodont apatite, already primary during the growth of conodont elements, represents a stable mineral phase which is commonly well preserved in sedimentary rocks during and after deposition. During diagenesis and metamorphism conodonts underwent systematic and irreversible changes in their chemical composition and structure leading to changes in optical properties. As a result of this transformation conodont elements are valuable thermal indices from deposition through at least low grade metamorphism. The application of conodont elements as palaeotemperature parameters has become a standard practice during the last decades and a large number of papers dealing with different aspects have been published. Thus, an overview on conodont colour alteration studies with respect to the composition of conodonts, surface texture, regional studies, hydrocarbon potential and contact metamorphism is given in the Supplementary material 1. 2.2.2. Colour Alteration Index (CAI) The conodont colour change during diagenesis has long been known (Ellison, 1944) but it has not been well explained and quantified until the work by Epstein et al. (1977) and later by Rejebian et al. (1987). This colour change is temperature-controlled and related to carbonization of small amounts of complex organic matter sealed within the conodont elements (Bustin et al., 1992; Marshall et al., 1999, 2001). Epstein et al. (1977) combined field collections with artificially heated samples to establish a semiquantitative scale of five colour intervals from 1 to 5 CAI, which can be assigned to different temperature regimes (50–300 °C) during diagenesis (Fig. 7). This index was extended by Rejebian et al. (1987) to higher temperatures (350–600 °C) and CAI values of 5–8. During artificial maturation of conodonts, progressive removal of carboxyl, hydroxy, alkoxy and carbonyl functional groups, loss of nitrogen containing moieties and development of highly ordered graphitelike carbon residues were observed (Marshall et al., 2001). According to these authors the removal of carbonyl functional groups as well as the loss of hydroxyl groups is combined with the transition from the interval CAI 1 to CAI 2. Further heating and colour change of conodont elements (CAI 3–CAI 6) are related to the principal removal of Corg

and the release of structurally bound water. These processes are irreversible. The CAI values are determined microscopically under incident light from thin edges or lighter parts of conodont elements by comparison with laboratory and/or field produced conodont standard sets produced by Harris (U.S. Geological Survey, Reston, Virginia). The standard consists of conodonts of variable size and morphology arranged in groups differing by half a CAI index from each other. Visual comparison with the CAI has a potential error of approximately half a CAI index (see chapter 3). In order to derive metamorphic temperatures for natural rocks, Epstein et al. (1977) and Rejebian et al. (1987) used an Arrhenius plot to extrapolate their experimental data to geologic time spans (Fig. 8). 2.2.3. Conodont colour vs. weathering, lithology, diagenetic alteration and tectonic stress It is still a matter of controversy whether different lithologies affect the colour of conodonts and there are some other pitfalls concerning the application of conodont colour. Epstein et al. (1977) used isolated conodonts from carbonates only and the effect of associated minerals on thermal alteration was not considered in their study. Other authors have shown that lithological differences might affect conodont colour. As one example, conodonts derived from highly bituminous carbonates generally show higher CAI values than those taken from other limestones of the same section (e.g. Königshof, 1992). Mayr et al. (1978) have shown that the CAI values of elements recovered from terrigenous clastic sedimentary rocks were higher than those taken from carbonates in an area which underwent the same thermal history. Other studies by Belka (1990) or Legall et al. (1981) confirmed and documented the influence of lithology on colour alteration of conodonts but they recorded differences that are within half a CAI unit only. Barham et al. (2012) have shown that conodont elements generally appear to have been less affected by diagenetic alteration even if they exhibit a high alteration index. Thus they are also very useful for oxygen isotope studies. The influence of dolomitization on conodont-bearing limestones has been recognized by Landing (1981) who described a conodont fauna in which the original brown–black colouration had been leached to yield white conodonts. Also other publications such as Helsen (1995),

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Time [h]

5 x 10 8 10 8 10 7 10 6 10 5

83

modification of high grade conodont CAI is presumably one of removal of organic material from the conodont structure. Finally, high pore pressure (confining pressure) may lead to CAI suppression by slowing down chemical transformation and/or in-mineral transport processes, e.g. in formerly overpressured rocks (e.g. Epstein et al., 1977; Hao et al., 2007).

10 4 10 3 10 2

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Fig. 8. Arrhenius plot after Epstein et al. (1977) and Rejebian et al. (1987).

Königshof (2003), Königshof and Werner (1994), March Benlloch and de Santisteban (1993), Nöth (1998), Nöth et al. (1991), Sanz-López and Blanco-Ferrera (2012), and Wardlaw and Harris (1984) have documented that diagenetic alteration influences the colour and the texture of conodont elements. Apatite pseudomorphs after calcite or dolomite, apatite overgrowth around conodont elements as a result of remobilization of apatite during dolomitization, and variable sizes of apatite crystals are common features. This type of alteration is very common and known since the 1970s from the Austrian Alps (e.g. Schönlaub and Zezula, 1975; Schönlaub et al., 1980), Japan (Kuwano, 1979), and Hungary (Kozur and Mock, 1977). This feature also occurs in reefal carbonates due to posttectonic dolomitization. Conodont elements obtained from dolostones show often a grey patina. According to Rejebian et al. (1987) this feature can be attributed to ascending hydrothermal fluids and/or oxidation of near surface organic matter. Thus, abnormally high CAI values which can occur in dolomitized limestones do not correspond to burial temperatures. Also oxidizing hydrothermal solutions can influence the colour of conodont elements (Harris, et al., 1990; Rejebian et al., 1987; Wardlaw and Harris, 1984). Some high indices of CAI 6 are not related to time and temperature conditions that produce similar values in anhydrous systems, but are a result of redox processes affecting conodonts at high temperatures for short periods of time (Rejebian et al., 1987). On the other hand, the organic material present within conodont elements seems to be less affected by weathering than dispersed organic matter of the host rock (e.g. Belka, 1993a; see Littke et al., 1991). Tectonic stress and strain can change colour of conodont elements but this effect has never been reported to increase CAI by more than one unit. Those conodont elements exhibit characteristic features such as fracturing and cleavage as a response to tectonic stress (Fig. 9). Cleavage is generally very common in conodonts with high CAI values (CAI 4–5.5) as it has been described by many authors (e.g. Epstein et al., 1977; Königshof, 2003; Kovács and Árkai, 1987; Kuwano, 1979; Raven and van der Pluijm, 1986). Cleavage also occurs at lower CAI values such as in the Variscan belt of Northern Spain, were CAI values of 3 exhibit cleavage structures (García-López et al., 1997). Washington and McCarthney (1982) recorded an apparent relationship between CAI and strain: increasing CAI values from a single limestone outcrop corresponding with increasing tectonic stress in the same outcrop. This observation has been confirmed in other studies such as in the Rhenish Massif, Germany, Montagne Noire, France and NE Hungary (Bender and Königshof, 1994; Sudar and Kovács, 2006; Wiederer et al., 2002). The mechanism involved in pressure solution

2.2.4. Application of conodont colour alteration studies A large number of regional studies have shown the potential of CAI in determining the transition from diagenesis to low-grade metamorphism in sedimentary basins as compiled in the Supplementary material 1. Conodont colour alteration indices have also been used in economic geology. The thermal window for commercial oil production generally corresponds with CAI 1.5 to 2.5. The gas window occurs within a range of CAI 2.5 to 4.5, while the cutoff for most hydrocarbon production is CAI 4.5 (Harris et al., 1980). In a number of studies CAI data have been used to assess thermal and burial history of sedimentary basins and in the exploration for hydrocarbons and ore deposits (e.g. Burrett, 1992; Harris et al., 1978; 1987; Hillier and Marshall, 1992; Nowlan and Barnes, 1987). Furthermore, conodont elements have also been used to determine thermal aureoles related to igneous rocks (e.g. Armstrong and Strens, 1987; Königshof, 1991; Nicoll, 1981; Swift, 1993; Wiederer et al., 2002). CAI data can indicate the extent and direction of buried intrusions, as described by Burnett (1987), Königshof (1991) and Nicoll (1981) among others. Surface texture and colour variations are the most characteristic features used to recognize contact metamorphosed conodont elements (e.g. Burnett et al., 1988; Königshof, 2003; Nöth, 1998; Rejebian et al., 1987; Voldman et al., 2008, 2009). Contact metamorphosed conodont elements differ considerably from texturally unaltered samples showing recrystallization and can contain a range of up to four CAI values within a single sample (Fig. 9). Burnett et al. (1988) and Königshof (1991, 1992) recorded a relationship between the growth of apatite crystals and CAI value: increasing crystal size corresponds to increasing CAI value in contact metamorphosed limestones. Recognition of contact metamorphosed conodonts provides relevant information about thermal anomalies and buried intrusions. 2.3. Graptolites 2.3.1. General remarks Taxonomic classification of the Pterobranchia has been considerably modified over the last years, mostly due to the application of cladistic approaches which led to the recognition of the modern pterobranch Rhabdopleura as an extant graptolite (Mitchell et al., 2013). The current classification subdivides the class Pterobranchia into two subclasses, the Cephalodiscida and the clonal Graptolithina (Maletz, 2014a). The former are “pseudo-colonial” pterobranchs, as the mature zooids are unconnected and thus, cannot be referred to as colonial in a strict sense (Maletz, 2014a). The Cephalodiscida comprise only a few mostly extant genera and some fossil genera. However, it is difficult to assign extinct taxa to this subclass as the separation of the mature zooids is hardly to prove in fossil material. With regard to organic petrology and thermal maturity studies, the Graptolithina (graptolites) are of much more importance. The zooids of the Graptolithina show a constant organic connection and thus are true colonial pterobranchs. The development of a graptolite colony begins with the sexually produced initial zooid which secrets the extracellular sicula. All post-sicular zooids are asexually evolved through budding and hence are clones of the initial zooid. The post-sicular zooids secrete the thecae and with the first theca the colony begins to grow throughout astogeny (Maletz, 2014c; Maletz et al., 2014). The only modern representative of the Graptolithina is the genus Rhabdopleura which includes several extant species. Apart from some families with uncertain taxonomic status, the Graptolithina enclose

Fig. 7. Naturally altered conodonts (CAI 1–CAI 6). Examples from the Rhenish Massif (Germany) and the Montagne Noire (France; after Königshof, 2003); listed temperature ranges after Epstein et al. (1977) and Rejebian et al. (1987).

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Fig. 9. A: Middle Devonian conodonts from a contact metamorphosed section (Odersbach quarry, Rhenish Massif, CAI 7). Note that the conodont surface looks smooth and unaltered, even if those sedimentary rocks have received temperatures of more than 800 °C. B: In a more detailed view contact metamorphosed conodonts often show recrystallization of apatite crystals. The size of the crystals is dependent on the heating event (Odersbach quarry, Rhenish Massif); SEM micrograph. C: Characteristic conodonts from contact-metamorphosed limestones (waterfall Romkerhall, Harz Mountains, Upper Devonian). Note the different CAI values (CAI 5, 6, 6.5 and 7) within one sample. D: Well preserved conodont element (Palmatolepis perlobata schindewolfi, CAI 3.5) from an Upper Devonian section (Donsbach quarry, Dillenburg, Rhenish Massif); SEM micrograph. E, F: Ductilely deformed conodonts (Pa-elements, Upper Devonian, Montagne Noire), CAI 5.5; SEM micrographs.

the Dendroidea (dendroids) and the Graptoloidea (graptoloids). Most dendroids are benthic and sessile but some taxa developed a secondarily derived planktic life style. The tubarium, the organic housing construction of the pterobranchs and formerly known as the rhabdosome in graptoloid research (Maletz et al., 2014), is erect, bushy or fan-shaped. All Graptoloidea are planktic. They have been recorded since the basal Tremadocian (Maletz, 2014a), while the dendroids had their first occurrence around the base of the Furongian, late Cambrian (Maletz, 2014b). The earliest graptolites, representatives

of the Rhabdopleuridae, appear in the Drumian, Cambrian Series 3 (Maletz, 2014b). With more than 600 genera and some thousands of species (Maletz, 2014a, c) the graptolites represent an important component of the early to mid-Palaeozoic marine macrozooplankton preserved in the fossil record. High abundance, distribution in a wide range of sedimentary facies types, complex construction of the colonies, rapid evolution, planktic life style of the graptoloids and some dendroids and also a lack of significant biogeographical differentiation except in the Ordovician


(Goldman et al., 2013), are all factors which particularly predestine the graptoloid graptolites as excellent biostratigraphical marker species. Ideally, the duration of a graptoloid biozone is as little as 300 000 years (Loydell, 2012), but most graptolite zone durations are longer. Sadler et al. (2009) estimated the mean duration at about 1.44 Ma for the Ordovician and 0.91 Ma for the Silurian. According to decay experiments with modern Rhabdopleura specimens, solely the tubarium of the graptolite animal, i.e. the organic housing construction, has a high preservation potential. Decay of the zooid soft tissue takes place within a few days and consequently fossil zooid remains are extremely rare (Briggs et al., 1995; Maletz, 2014b, c). Thus, only the tubaria will be encountered in polished blocks used in organic petrological studies. Graptolites are especially abundant in marine, carbonaceous black shales, also described as graptolite shales. In this facies graptolites are usually preserved as flattened organic films (Maletz, 2014c), but may be preserved in three dimensions filled with early diagenetic pyrite. Graptolites are also found with less abundance in sandstones, marls, limestones and authigenic cherts from the inner shelf through to the ocean floor. In the latter two rock types preservation can be three-dimensional and exceptionally good. The graptolitic black shales were probably deposited in deep water on the outer shelf, continental slope and ocean floor in marginal upwelling regions with increased nutrient supply and enhanced bioproductivity (Cooper et al., 2012). At least during the Ordovician, many graptoloids inhabited an offshore deep water biotope below the epipelagic zone (water depth N200 m) of the outer shelf, continental slope and ocean floor with dysaerobic water, but the preferred habitat was probably the epipelagic zone with aerobic water from the inner shelf through to the ocean floor (Cooper et al., 2012). Reconstructions of the graptolite animal and its life style are discussed in detail in Maletz (2014c). 2.3.2. Graptolite morphology under reflected white light in polished blocks In pre-Devonian or in marine Palaeozoic sedimentary rocks where vitrinite is absent or rare, the reflectance of graptolite fragments serves as one of the major maturity parameters. In practice, Ordovician to Lower Devonian rocks are the main field of application of graptolites in maturity studies (Supplementary material 2). Usually, graptolite fragments can easily be identified in polished blocks and reflected light microscopy as they are clearly anisotropic. As pointed out by Petersen et al. (2013) vitrinite-like particles recorded in Lower Palaeozoic sedimentary rocks, lacking the diagnostic graptolite features and exhibiting a more varied morphology are probably remnants of graptolites. Larger graptolite pieces are mostly thin elongated, lathshaped bodies and may show the tubarium with parts of the thecae or the tubarium wall enclosing the common canal (Goodarzi, 1984a; Hoffknecht, 1991) which connects the individual thecae (Maletz et al., 2014). In addition, graptolite fragments may portray distinct morphological features such as a finely lamellar structure which can be best recognized under crossed polars. This structure has been described and figured in detail by Clausen and Teichmüller (1982), Hoffknecht (1991) and Teichmüller (1978). Basically, the tubarium wall is composed of an inner and outer organic layer. The inner layer is the fusellar tissue, the outer layer is the cortical tissue which in turn consists of an alternation of cortical and sheet fabric. Ultrastructural studies show that both layers are formed by fibrils which vary between 20 and 400 nm in diameter (e.g. Bates et al., 2009). The lamellar structure can be assigned to the cortical tissue and shows light and dark coloured lamellae. Under high magnification the fibrous structure caused by the fibrils is discernible (Teichmüller, 1978, Figs. 2b, 3b). The lamellar structure is more pronounced in slightly weathered samples of low thermal alteration. The light coloured lamellae (cortical fabric) are 2–5 μm wide and show bireflectance and reflectance pleochroism under polarized light, while the dark coloured lamellae (sheet fabric) are b 1–1.5 μm wide and do not change in colour. The proportion and also the width of light and dark coloured lamellae do influence the reflectance values (Hoffknecht, 1991; Teichmüller, 1978). Teichmüller (1978) did not

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observe remnants of the fusellar tissue in her highly coalified material (GRmax 5.1–10.0%). She assumed that due to the high thermal alteration, the original rather loose and porous tissue was destroyed without leaving any recognizable traces. In samples from the Road River Group (GRmax 2.66–5.74%, following Link et al., 1990) Goodarzi et al. (1992) identified graptolite fragments with the fusellar layer preserved and also Hoffknecht (1991) noted that fusellar remains have rarely been encountered, most likely due to the compaction of the porous tissue. Under reflected white light graptolite fragments exhibit two types of surface morphology which Goodarzi (1984a) circumscribed as granular and non-granular. The granular surface has a fine granular to reticular texture and the granular fragments are softer, lower reflecting and show weaker anisotropy. Non-granular graptolite fragments that often show fine structural details (e.g. light and dark coloured lamellae), are hard, brittle, high reflecting and show stronger anisotropy (e.g. Goodarzi, 1984a; Goodarzi and Norford, 1985, 1987; Goodarzi et al., 1992). Riediger et al. (1989, Fig. 6) distinguished two nongranular varieties, an angular, blocky shaped variety and a lath shaped one which also exhibits a higher reflectance and stronger bireflectance. It is not clear if the predominance of either type relies on the petrography of the host rock. While e.g. Goodarzi and Norford (1985) and Goodarzi et al. (1992) reported the non-granular type as being more abundant in shales and the granular type more common in carbonate matrices, Hoffknecht (1991) could not prove such a reliance, even though his study is based on an extraordinarily large sample set. Various interpretations have been suggested to explain the nature of the granular and non-granular texture. The two types may represent different parts or tissues of the tubarium (Goodarzi, 1984a; Goodarzi et al., 1992; Link et al., 1990). Riediger et al. (1989) assumed that the granular type represents the remains of the soft-bodied animal which is, however, very unlikely because the zooids decay within days and readily before burial (Briggs et al., 1995). Bacterial decomposition after the death of the graptolite animal or very early after burial may also explain the occurrence of two different textures in graptolite fragments (Hoffknecht, 1991). 2.3.3. Graptolite reflectance, bireflectance and colour change during diagenesis Graptolite remains are occasionally encountered in palynological preparations of Ordovician to Devonian sedimentary rocks (e.g. Batten, 1996a, p. 1035, Plate 4; Traverse, 2008, Plate 18.3e; Tyson, 1995, Fig. 20.1) but, if at all, due to the processing procedure only small fragments will be preserved. If found dispersed in palynological slides they are classified as zooclasts. Progressive changes in graptolite colouration due to increasing palaeotemperature have never been investigated in detail. Dorning (1986) reports that the colour of graptolite fragments in transmitted light ranges from various shades of brown to black with increasing maturity and often differs only slightly from that of chitinozoans and scolecodonts. There seems to be some potential for using colour alteration of graptolites for assessing thermal maturity, yet a colour scale similar to that applied for some palynomorph groups has not been proposed. Graptolite reflectance has been used in numerous studies focussing on regional maturity trends including differentiation of metamorphic subfacies, hydrocarbon and ore exploration and contact metamorphism (for references see Supplementary material 2). In reflected light graptolites show strong anisotropy which increases with higher thermal maturity levels. Other optical parameters such as reflectance, refractive index and absorptive index follow this trend. Obviously, bireflectance depends to a certain extent on rock matrix and orientation of the polished block to the bedding. Graptolites preserved in limestone or marls have a lower bireflectance than those embedded in shales. Suppression of graptolite reflectance has also been reported from oil shales (Goodarzi et al., 1992). Anisotropy is more pronounced in sections perpendicular to bedding than parallel to bedding. Link et al. (1990) proposed maximum reflectance of graptolites (GRmax) as the most

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reliable parameter. Differences also exist between reflectance, anisotropy and bireflectance of non-granular and granular graptolite fragments, the latter exhibiting lower values in all three parameters (Goodarzi, 1984a; Goodarzi and Norford, 1985; Goodarzi et al., 1992). Graptolite reflectance may also be affected by oxygen level in the water column and in the sediment. Cole (1994) reported lower reflectance values in graptolites deposited under anoxic conditions compared to those from oxic environments. Another factor influencing the optical properties of graptolites is weathering. Goodarzi et al. (1992) observed highly reflective oxidation rims and cracks while Xiaofeng et al. (1993) showed that weathering reduces reflectance values. Petersen et al. (2013) published a large set of data on graptolite reflectance which they compared to the reflectance of zoomorphs (chitinozoans), vitrinite-like particles and solid bitumen. Graptolite reflectance increases in a similar way as humic organic matter, but at a higher rate. They provided an equation for conversion of graptolite reflectance (GRr) into vitrinite reflectance: VRc ¼ 0:73 GRr þ 0:16

ð5Þ

carefully explaining the role of vitrinite-like particles and how the correct, low reflecting population should be selected. Petersen et al. (2013) also provided a correlation of graptolite reflectance and CAI: CAI ¼ 1:31 GRr þ 0:49:

ð6Þ

In addition, equations to convert graptolite reflectance into vitrinite reflectance and reflectance of zoomorphs (chitinozoans, scolecodonts) and zooclasts (hydroids) have been published by e.g. Bertrand (1990a, 1993), Bertrand and Héroux (1987), Bertrand and Malo (2001, 2012) and Yang and Hesse (1993). 2.3.4. Chemical alteration of graptolites during diagenesis The relatively durable organic tubarium (= rhabdosome) is secreted by the zooids and has a high preservation potential in dysoxic and anoxic environments. Even under oxic conditions preservation is possible, provided that the graptolite remains are quickly embedded in the sediment (Maletz, 2014b). Using amino acid analyses and histochemical tests combined with transmission electron microscopy, the chemical composition of the tubaria of a modern Cephalodiscida and graptolites (Rhabdopleura and fossil material) has been investigated since the 1960s when e.g. Jeuniaux (1963) excluded the presence of chitin in Rhabdopleura normani and Manskaâ and Drozdova (1962) as well as Foucart et al. (1965) reported various amino acids in the hydrolysates of fossil graptolites. However, the presence of amino acids in graptolites could not be confirmed in later studies (for a review of these early publications see Briggs et al., 1995; Manskaya and Drozdova, 1968; Urbanek, 1976). Because of high amounts of glycine, serine and alanine a scleroproteic material was assumed. This conclusion is still widely accepted in graptolite publications (e.g. Maletz, 2014c). More recent approaches to decipher the composition of modern and fossil pterobranchs applied FTIR and μ-FTIR spectroscopy (e.g. Bustin et al., 1989; Ganz et al., 1990; Hoffknecht, 1991; Suchý et al., 2004), μ-Raman spectroscopy (Suchý et al., 2004), Curie Point-pyrolysis–gas chromatography/mass spectrometry (e.g. Briggs et al., 1995), TMAH assisted thermochemolysis–gas chromatography/mass spectrometry (Gupta, 2014; Gupta and Briggs, 2011; Gupta et al., 2006) and timeof-flight secondary ion mass spectrometry (Liu et al., 1995). Some of these studies pointed to a chemical composition of fossil graptolites that is dominated by aromatic and aliphatic moieties. These observations supported by FTIR spectroscopical (Bustin et al., 1989) as well as by pyrolytical analyses (Briggs et al., 1995; Gupta et al., 2006) indicate that the fossil organic material in graptolites did not derive from proteinaceous material and, therefore, a selective preservation of the scleroproteic matter can be excluded. Alternatively, it has been

speculated that the fossil material in graptolites is formed during diagenesis either by in situ polymerization of lipids (Gupta, 2014; Gupta et al., 2007) or a replacement by external aliphatic geopolymers (Briggs et al., 1995). However, detailed information about the exact thermal maturity of the investigated graptolites is often lacking and therefore knowledge about the process of diagenetic formation of fossil graptolite material is very limited. Bustin et al. (1989) performed maturation simulation experiments applied to graptolites. With increasing thermal stress they observed a decrease of aliphatic moieties accompanied by an increase in aromaticity including ongoing condensation of aromatic systems. 2.4. Foraminifera 2.4.1. General remarks Foraminifera are single-celled Protozoa ranging from the Early Cambrian to Holocene. They are excellent biostratigraphic markers in marine successions as they are widespread (from the deep sea to marshes), abundant (even rock-forming) and have some rapidly evolving lineages. They occur in brackish and also rarely in freshwater environments but the highest diversity is reached in normal marine waters (Armstrong and Brasier, 2005). Foraminifera are primarily benthic but the tests of planktic species considerably contribute to the accumulation of pelagic sediment (calcareous ooze). The majority of foraminifera tests are composed of calcium carbonate with thin inner or outer organic linings. In addition, organic material may be present as lining of pores and between calcite layers (McNeil et al., 1996). In agglutinated foraminifera the tests are made up of detrital grains and organic particles cemented by various substances, e.g. calcite/aragonite (Roberts and Murray, 1995) or an organic matrix (Allen et al., 1999, 2000; Bender, 1989). The ultrastructure of the organic matrix in agglutinated foraminifera has been studied in detail and various morphotypes can be distinguished. Besides the organic cement, the grains in agglutinated foraminifera are coated with an organic envelope and an inner and outer organic lining may be developed (Bender, 1989; Bender and Hemleben, 1988). A rarely studied foraminifera group, the ‘allogromiids’, have an organic or predominantly organic test wall (Gooday, 2002; for a review on the various types of foraminifera test walls see Ehrlich, 2010). Preservation of foraminifera test walls is affected by numerous diagenetic processes such as decay of the organic matrix, thermal maturation, hydrocarbon impregnation, mineralization or dissolution (e.g. McNeil, 1997; McNeil et al., 2010). The acid-resistant, organic linings of calcareous and agglutinated foraminifera frequently occur in palynological samples from marine and slightly brackish water environments (Stancliffe, 1996). Very high abundances are reported from upwelling areas (Tyson, 1995, p. 336). These palynomorphs have also been referred to, among others, as microforaminifera, palynoforaminifera and Scytinascia, but microforaminiferal lining is the only appropriate term (Fechner, 1999; Stancliffe, 1989, 1996). These linings are small (usually less than 150 μm in diameter) and thus it was thought that they represent an internal organic layer of very small foraminifera (‘microforaminifera’). However, Fechner (1999) and Stancliffe (1989) demonstrated that in microforaminiferal linings the terminal chambers are frequently missing, which is not surprising as the lining is the thickest in the proloculus and progressively thins out towards ontogenetically younger chambers. The thin-walled linings of the terminal chambers are more prone to microbial degradation as well as chemical decomposition and are easily detached (e.g. during transport or when harsh processing techniques are applied). Therefore, the microforaminiferal linings in palynological preparations are expected to be smaller than the foraminifera test they are derived from and must not inevitably be assigned to very small foraminifera. The earliest microforaminiferal linings have been described from the Lower Cambrian Lontova Formation, Russia (Winchester-Seeto and McIlroy, 2006). Although microforaminiferal linings are regular constituents in palynological preparations of marine


sediments, they have very rarely been found in thin sections or polished blocks. Mišík and Soták (1998) report organic linings of small foraminifera in thin sections of Mesozoic limestones and cherts. To our knowledge microforaminiferal linings have not been studied in polished blocks.

to a vitrinite reflectance of 0.5% (Gallagher et al., 2004). A consistent and progressive increase of FCI values occurs up until FCI 6, whereas an increase of vitrinite reflectance is less marked in this range (Fig. 10). Thus, FCI is a suitable measure in the immature and mature stages. From FCI 6 (very dark grey to very dark greyish brown) onwards, the colour change is less pronounced but vitrinite reflectance increases rapidly. The highest FCI value (FCI 10, black opaque) is reached at a vitrinite reflectance of 3.5% and higher (Fig. 10). At very high thermal maturity black foraminifera become bleached white (Rosen, 1991) but this stage is not covered by a FCI category. In several experiments, both recent and fossil agglutinated and calcareous foraminifera tests were subjected to increasing temperatures and pressures using various host sediments (Alabusheva, 1990; Hagerman Reed, 1999; Kontrovitz and De Hon, 1983; McNeil et al., 1996; Rosen, 1991). None of the experiments in which agglutinated foraminifers were included were as profound as those conducted by Kontrovitz et al. (1992) on ostracod valves (see below). Some results of the foraminifera test experiments are difficult to interpret and partly contradictory. Discolouration of agglutinated tests is very different from calcareous ones, i.e. agglutinated foraminifera exhibit a distinct colour change upon heating while colour alteration in calcareous tests with little or no organic cement is either completely lacking or negligible and does not correlate with temperature (Hagerman Reed, 1999; McNeil et al., 1996). McNeil et al. (1996) showed that heating recent and fossil

2.4.2. Colour alteration and mineralogical changes of foraminifera during diagenesis During thermal maturation the organic cement of agglutinated foraminifera shows a distinct colour change from white to black opaque (Fig. 10). Also microforaminiferal linings undergo a marked colouration change. In addition, agglutinated and calcareous foraminifera experience mineralogical alteration with increasing burial diagenesis. Discolouration of foraminifera has been used to develop a 5-point scale (Alabusheva, 1990), a 6-point scale for agglutinated foraminifera (Rosen, 1991) and a 10-point scale; the latter established by McNeil et al. (1996) is also based on agglutinated foraminifera and named the Foraminiferal Colouration Index (FCI). Textural and mineralogical changes of the foraminifera tests allow to distinguish four zones (A–D) of burial diagenesis (McNeil et al., 1996). Agglutinated foraminifera from low maturity sediments show white to light grey colours (FCI 1 and 2). With increasing maturation level the colour of agglutinated foraminifera specimens darkens from light brownish grey to grey (FCI 3), grey to greyish brown (FCI 4) and dark grey to dark greyish brown (FCI 5). FCI 3.5 is approximately equivalent

Zone A

87

Zone C

Zone B

Zone D

Vitrinite Reflectance VR r [%]

4

3

data from McNeil et al. (1996) data from Gallagher et al. (2004) 2

1

0 0

1

2

3

4

5

6

7

8

9

10

Foraminiferal Colouration Index (FCI)

Fig. 10. Zones of burial diagenesis (A–D), Foraminiferal Colouration Index (FCI) of McNeil et al. (1996, Fig. 8) and correlation between vitrinite reflectance and Foraminiferal Colouration Index. Data from McNeil et al. (1996, Fig. 9–11) and from Gallagher et al. (2004, Fig. 12). Each datum point of Gallagher et al. (2004) is based on numerous samples.

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foraminifera tests in the presence of air resulted in bleaching of the specimens which is probably due to oxidation and subsequent volatilization of the organic material. Under inert conditions recent specimens with large amounts of organic cement darkened significantly while those with little or no cement darkened slightly or not at all. Specimens which were soaked in crude oil prior to heating also turned darker especially in areas where pores are filled with clay or were clay adheres to the test (McNeil et al., 1996). In this respect these experiments showed different results from those with ostracods (see below) as conducted by Kontrovitz et al. (1992). In most cases, foraminifera colour change during thermal maturation does not seem to be achieved by impregnation of the test through mobilized hydrocarbons. Instead, darkening of the organic cement gives rise to the discolouration of the agglutinated test wall. Rare exceptions are agglutinated foraminifera from a hydrocarbon seepage system reported by McNeil et al. (2010). These tests are impregnated by hydrocarbons, show silicification differing from normal thermal alteration trends, a wide colour range and are darker than those tests from the same depth interval which have been affected by burial diagenesis only. While in agglutinated foraminifera colour is thought to be only slightly or not at all influenced by the surrounding sedimentary rock the amount of organic cement in the test wall does have an effect on discolouration, i.e. test walls with high amounts of organic carbon display darker colours than those with little or no organic cement (McNeil et al., 1996; Rosen, 1991). When using the FCI it must be taken into consideration that in agglutinated foraminifera the amount of organic material varies from one genus to another influencing colour and colour alteration to a certain extent. Ideally, assemblages of similar taxonomic composition should be compared. Apart from the quantity of organic cement, pore overpressure conditions may also affect the colour of foraminifera tests, i.e. colour change can be retarded or bleaching may even occur in overpressured zones (Issler et al., 2002; McNeil, 1997; McNeil et al., 1996). The interpretation of colour change in agglutinated foraminifera with regard to thermal maturity is further complicated by the fact that humic acids may cause differential post-mortem staining of the tests (Wilson and Vincent, 2014). FCI measurements are a very fast, inexpensive and easy tool used to estimate thermal maturity. Processing for foraminifera does not involve employment of highly corrosive and very toxic substances as is routinely necessary in palynological preparations. Consequently, FCI can be utilized at well sites to provide preliminary evidence for thermal maturity. Depending on the rock type, standard processing techniques for separating foraminifera involve chemicals such as tensides, organic solvents, sodium salts, domestic bleach (sodium hypochlorite) or hydrogen peroxide. The use of oxidants should be avoided as they may alter foraminifera colouration (McNeil et al., 2000). The possible influence of other chemicals on the FCI has not been studied systematically. To preclude any undesirable colour change, non-chemical disaggregation techniques are preferable. Unconsolidated or indurated sediment can be broken down by using e.g. warm water, repeated freeze-thawing, liquid nitrogen or a combination of these and other physical extraction procedures. FCI values can easily be determined under a dissecting binocular microscope by comparing overall foraminiferal colours with the Munsell Soil Colour Chart. Colour variation in an assemblage is documented by analysing some tens of specimens resulting in a histogram and by calculating average FCI values and standard deviation (Gallagher et al., 2004; McNeil et al., 1996). Microforaminiferal linings change in colour from pale orange (Fig. 11) to red brown, brown and black when subjected to increasing thermal maturity and seem to react similarly to other thick-walled palynomorphs (Stancliffe, 1996). However, a distinct colour change does not appear in the immature to early mature stage (Fig. 11). Discolouration of the inner organic lining of agglutinated foraminifera is probably also similar to that of their organic cement but the inner lining is much darker than the overall colour of the test wall (McNeil et al., 1996). An index based on the colouration change of microforaminiferal linings has not yet been proposed. As the wall of the proloculus is thicker and therefore darker

in colour than that of succeeding chambers (Stancliffe, 1996), colour determination should be done on the ontogenetically oldest chambers. An estimation of thermal maturity using microforaminiferal linings is further complicated by the fact that the initial colour of the linings may vary, as demonstrated by Winchester-Seeto and McIlroy (2006) who observed dark brown to black linings but no blackened acritarchs in the same sample. Globigerina bulloides, a planktic calcareous foraminifera, was used by Hagerman Reed (1999) to conduct heating experiments in a closed stainless steel container filled with seawater. In specimens that were heated up to 200 °C the organic matter did not display any colour change whereas under the very same conditions modern Pinus pollen showed discolouration. In open crucibles and at higher temperatures (300–500 °C) the organic matter of the foraminifera darkened progressively from yellow to brown. Hence, organic matter of G. bulloides starts to change colour at higher temperatures compared to modern pollen. Apart from thermally induced colour changes, both agglutinated and calcareous foraminifera tests respond to increasing burial diagenesis by progressive mineralogical changes (e.g. Brunner, 1994). However, not only burial temperature and burial history can influence diagenetic overprinting of foraminiferal tests but also various other factors such as water depth and temperature, sediment composition or pore water chemistry (e.g. Pearson and Burgess, 2008). The transformation of aragonite to calcite in primary aragonitic foraminifera is known to occur within the upper part of the oil window (Reiser, 1988). McNeil et al. (1996) found mineralogical and textural changes with increasing thermal maturity in agglutinated and calcareous foraminifera which were then assigned to four zones (Zone A–D, see Fig. 10). Zone A is characterized by the lack of mineralogical changes but a reduction of organic cement and porosity. In Zone B the organic cement is further diminished and due to newly formed quartz crystals and partial secondary silica overgrowth on detrital quartz grains the porosity is also reduced. Zone C shows well silicified, glassy translucent tests with much reduced porosity due to quartz, kaolin and smectite precipitation. However, much of the original grain structure is still preserved. In Zone D the primary porosity is more or less lost and the original grain texture destroyed due to silicification. To date, the FCI and the burial diagenetic zones of McNeil et al. (1996) have found application as an exploration tool in only few surveys (Supplementary material 3). However, the scales have been used to decipher thermal alteration of sediments in very different tectonic and environmental settings such as intrusion-related hydrothermal gold mineralizations (Gunson et al., 2000), diamond-bearing kimberlite deposits (McNeil et al., 2000), hydrocarbon seepage systems (McNeil et al., 2010) and palaeo-overpressure zones in sedimentary basins (e.g. Issler et al., 2002; McNeil, 1997). Autofluorescence of the organic cement in agglutinated foraminifera was investigated in a preliminary study by Stasiuk and McNeil (2000). Only three samples were considered covering a vitrinite reflectance range from 0.15 to 0.56% VRr. Foraminifera with the lowest thermal maturity show the most intense autofluorescence with a λmax of 455–480 nm and a red/green quotient of 0.14–0.21. With increasing maturity both parameters increase up to λmax of 630–650 nm and a red/green quotient of 1.30–1.60. Hence, fluorescence colour changes from blue–green in the lowest maturity level (0.15–0.20% VRr) over orange–yellow (0.30–0.35% VRr) to brown–yellow to orange–brown in the highest maturity level (0.56% VRr), i.e. increasing maturity corresponds to a red shift of fluorescence colour. Surprisingly, in this low maturity range the red shift in agglutinated foraminifera seems to be more rapid than in miospores and alginites (Stasiuk and McNeil, 2000). 2.4.3. Chemical composition of organic matter in foraminifera Studies on the geochemistry of calcareous foraminifera tests focussed on oxygen isotope composition and the Mg/Ca ratio which are important palaeotemperature proxies especially when applied in combination. Much less is known about the chemical composition of the inner and outer organic linings and the organic cement in


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Fig. 11. Microforaminiferal linings from the Lower Jurassic, Hils Syncline, Germany. Microforaminiferal linings show distinct colour variation in the respective specimens and no pronounced colour alteration in the immature stage due to thermal maturation. (a) and (b) borehole Wenzen, 55.90 m; (c) borehole Dielmissen, 77.60 m; (d) borehole Dohnsen, 82.45 m. For more information on this borehole programme in the Hils Syncline see al Sandouk-Lincke et al. (2013).

agglutinated foraminifera. Following Langer (1992) glycosaminoglycans are highly relevant constituents of organic cements and linings of foraminifera. Glycosaminoglycans are polysaccharides consisting of aminosugars and glucionic acids appearing in organisms attached to proteins as so-called proteoglycans. Therefore, μ-FTIR spectroscopy and pyrolysis–gas chromatography/mass spectrometry analysis of organic material of foraminifera revealed that proteinaceous components with structural similarity to collagen are also dominant while carbohydrates are present only in minor amounts (Allen et al., 2000). Weiner and Erez (1984) identified significant amounts of various proteins especially in the soluble fraction obtained after decalcification of calcareous foraminifera with EDTA, yet they pointed also to a dominance of glycosaminoglycans in the insoluble fraction remained after EDTA treatment. However, with regard to the secondary structure of the protein component, the chemical composition of organic cements and linings differs slightly between genera. The organic preservation of fossil foraminifera is even less well known. Microforaminiferal linings are frequently regarded as chitinous, “chitinoid”, pseudochitinous or “tectinous” (e.g. Banner et al., 1973; Traverse, 2008) although their organic composition was never elucidated in a detailed study. With increasing thermal maturity μ-Raman spectroscopy of the organic cement in agglutinated foraminifera shows a progressive, but non-linear, increase in the intensity of the G and D bands and a decrease in the band width relative to the intensity.

Above the late mature oil window the G and D bands are very similar to those of graphite (McNeil et al., 2013). 2.5. Ostracods 2.5.1. General remarks Ostracods are commonly encountered in micropalaeontological samples. The bivalved shell (carapace) consists of chitin or low-magnesium calcite. Ostracods inhabit marine, brackish and freshwater environments and their fossil record dates back to at least the Late Ordovician (Siveter et al., 2014) or even late Cambrian (Harvey et al., 2012). Due to their high fossilization potential and diversity they are widely used in biostratigraphy and in palaeoecological studies. Exceptional circumstances such as phosphatization or pyritization may allow the preservation of ostracod soft parts (e.g. Siveter et al., 2014). Organic linings of mineralized carapaces have rarely been reported from palynological preparations (Braun, 1997; Goodall et al., 1992; Tyson, 1995, 212). These linings have also been isolated from single ostracod valves as old as Silurian (Gocht and Goerlich, 1958; Sohn, 1958). During diagenesis carapaces composed of low-magnesium calcite experience recrystallization, the degree of which being dependent on host sediment lithology and thermal maturation. At early and shallow diagenesis the original carapace is replaced by neomorphic calcite. At

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higher diagenetic alteration replacement and void fills with pyrite, ferroan calcite, ferroan dolomite, siderite and finally sphalerite and barite occur (Bennett et al., 2011). Simultaneously, the carapaces show distinct surface features such as pits, tiny knobs, short ridges, and striations. Following experimental studies in which ostracod shells were heated up to 360 °C under varying pressures, these surface characteristics are diagnostic and can be clearly related to temperature conditions although sediment composition, pore water chemistry and to a lesser extent also pressure may retard or accelerate the reaction (Ainsworth et al., 1990; Kontrovitz, 1987; Kontrovitz and Slack, 1995; Kontrovitz et al., 1983). 2.5.2. Colour alteration and mineralogical changes of ostracods during diagenesis Besides microscopic surface modifications, recrystallization and replacement by a range of diagenetic minerals, ostracod valves undergo a progressive colour change during thermal maturation (Fig. 12). Kontrovitz et al. (1992) simulated this colour alteration by subjecting modern ostracods to various temperatures as well as pressures and using organic matter-free to organic matter-rich sediments (12.8% organics) and crude oil. When modern unaltered ostracods were heated up to 650 °C without adding any sediment they slightly changed colour becoming first grey and subsequently white again, probably because organic material was broken down. In organic matter-free sediment no

pronounced discolouration was provoked. Sediment containing siderite or hematite caused insignificant colour alteration with an inconsistent trend as dark brown to black colours frequently encountered in fossil assemblages were not attained. However, when the ostracods were embedded in organic matter-rich sediment during the experiments the shells turned brown, dark grey and finally black. Temperatures of less than 300 °C were required to induce this colour change. Slight colour differences were noticed depending on the presence of soft parts in the shells. Specimens heated in organic matter-lean sediment to which crude oil was added developed grey, olive and brown colours over time even when the temperature did not exceed 50 °C. Hence, only ostracod shells which were mixed with organic matter-rich sediment or crude oil upon heating showed a similar colour path as fossil ostracods. Experimental reheating of black or dark greyish brown fossil ostracods resulted in lighter colours at about 400 °C. Specimens subjected to temperatures of 900 °C turned white. Thus, at temperatures above 400 °C the colouring agent starts to decompose. However, even near thick igneous intrusions and in metamorphic rocks ostracod shells display dark grey to black colour (Ainsworth et al., 1990; Kontrovitz et al., 1992). The experimental study of Kontrovitz et al. (1992) clearly demonstrates that the colouring is not or only to a small extent due to indigenous organic material of the ostracod shells (e.g. organic linings, organic sheets within the shell) but that organic compounds from the surrounding sediment are absorbed during heating. Since proteinaceous

Hydrocarbon Maturity Stages Immature

Mature

Overmature

Postmature

7

Ostracod colour (6,5)

6 (5,5)

5 (4,5)

4 (3,5)

3 (2,5)

2 (1,5)

1 0.2

0.3

0.4

0.5

0.6 0.7 0.8 0.9 1.0 1.2

1.4 1.6 1.8 2.0

2.5

3.0

Vitrinite Reflectance VR r [%]

Fig. 12. Correlation between vitrinite reflectance and ostracod colour indices of Ainsworth et al. (1990, Plate 1 and Fig. 2). The seven whole number levels correspond to Munsell Soil Color Chart chips. Half units are calculated mean ostracod colour indices based on the complete ostracod assemblage.


and chitinous material in and covering the ostracod carapace is usually decomposed during very early diagenesis, organic compounds such as hydrocarbons can easily invade the permeable shell (Ainsworth et al., 1990). Based on numerous samples from the Irish and UK offshore, ostracod colour change has been described in detail by Ainsworth et al. (1990). These authors also suggest a 7-stage colour chart where the lowest alteration level comprises colourless/white valves (level 1). Usually, this level corresponds to Holocene and unaltered Cenozoic/Mesozoic specimens. With increasing thermal maturity ostracod colour passes through yellow, orange, amber, red–brown and black (level 7). Ainsworth et al. (1990) published a broad correlation between vitrinite reflectance and ostracod colour (Fig. 12). Yellow to brown ostracod shells are characteristic for the immature as well as the mature stage and black colours prevail above a vitrinite reflectance of 1.7%. The most pronounced colour shift occurs in the late immature and mature stages, hence, this parameter is ideal for maturity studies related to oil generation and exploration. The application of the ostracod maturity parameter has been demonstrated by Ainsworth et al. (1990) who studied 34 exploration wells in offshore Ireland and Western UK. The gradual alteration in ostracod colour is burial depth-related, i.e. can be attributed to the overall palaeogeothermal gradient while rapid colour change is due to local heating by igneous intrusions (see Supplementary material 3). Furthermore, ostracod colour can easily be used to trace reworking from older to younger sedimentary rocks provided that they differ with regard to maturity levels. As has been demonstrated by e.g. Ainsworth (1987), Ainsworth et al. (1990), Kontrovitz and Slack (1995) and Kontrovitz et al. (1992) ostracod discolouration is related to thermal alteration and can be used as a rough organic maturation parameter (Supplementary material 3). Nonetheless, some shortcomings must be considered when using ostracod colour indices. As the colour change is attributed to organic material invaded from the sediment into the carapace, the amount of organic carbon of the surrounding rock controls colour alteration at least to some extent. This is especially important in organic matter-lean sedimentary rocks such as white to light grey limestones and chalks. From a TOC of 0.2–0.5% up an even higher organic carbon content does not seem to further influence the ostracod colour (Ainsworth et al., 1990). The nature of the organic matter in sedimentary rocks probably also affects ostracod colour. It can be expected that source rocks containing Type I and II kerogen will stain the shells more noticeably than source rocks with Type III kerogen provided that thermal maturation is in or above the oil window. However, this aspect has not yet been systematically studied. Ideally, ostracod assemblages from similar rock types should be used in thermal alteration studies. Mineral impregnations may also be a minor contribution to colour change. This is especially true for pyritization and oxidation of pyrite. Finally, processing techniques must not be neglected. Unconsolidated sediments and slightly indurated sedimentary rocks are routinely disaggregated by using hydrogen peroxide or sodium hypochlorite solution which are both strong bleaching agents and may alter ostracod colour. Acid processing (e.g. acetic acid, formic acid) usually applied to limestones can change ostracod colour considerably, i.e. from dark brown or black to white even before damage by acid etching is discernible. Curiously, assemblages from weathered limestones seem to be less prone to processing induced colour alteration than those from unweathered samples (Tarsilli and Warne, 1997). To which extent other processing techniques (e.g. Glauber's salt, tensids/surfactants such as Rewoquat®, kerosene) may influence ostracod colour has not been systematically studied. There is no detailed study on the macromolecular composition of the organic remains of fossil ostracods available and the chemical alteration of the organic matter in ostracod carapaces during thermal maturation has never been elucidated. By analogy with modern ostracods the organic linings and soft parts remaining after decalcification of ostracod valves are believed to be chitinous or to consist of chitin and/or protein (Sohn, 1958). However, no up-to-date study was devoted to organic geochemical analyses of these ostracod residues.

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2.6. Conchostracans 2.6.1. General remarks Conchostracan carapace valves are rarely encountered in organic petrological preparations (McDonald, 2007) or dispersed in palynological slides (Batten, 1996a, p. 1035; Lana and Carvalho, 2002). Nevertheless, they can be very abundant in thin horizons reflecting harsh environmental conditions (Webb, 1979) and as the valves are composed of organic material they show a colour change with increasing thermal maturity (Tasch, 1982). In maceral and palynofacies analyses conchostracan valves would be classified as zooclasts, zoobenthos or faunal relics (Tyson, 1995, p. 344). The paraphyletic conchostracans (clam-shrimps) are now regarded as a provisional collective name for several suborders of the Diplostraca, an order of the Crustacea. The fossil record of conchostracans dates back to the Early Devonian. Extant and fossil conchostracans are nonmarine except for some Devonian and Mississippian forms which probably inhabited deltaic and estuarine brackish water environments and perhaps may have even migrated into shallow marine waters. However, from the Pennsylvanian to Holocene they were restricted to freshwater habitats such as lakes, inland ponds and ephemeral water bodies. 2.6.2. Colour alteration and mineralogical changes of conchostracans during diagenesis In a preliminary study Tasch (1982) observed baked conchostracan valves and valves of different colours and degrees of deformation in sedimentary rocks of Jurassic age intercalated between two basalt flows. Modification of the valves varied from a brown crust-like exterior to a white or blue enemeloid appearance. In heating experiments Tasch (1982) distinguished seven valve colour classes (“valve geothermometry”) some of which are also characterized by the deformation of the valves depending on the applied temperature range. In these experiments the valves remained unchanged when temperatures did not exceed 300 °C. At higher temperatures the valves were first charred, then displayed a surficial light-brown crust as well as cracking and curling and later the colour changed to bluish–green. Increasing temperatures resulted in a white enameloid or porcelaneous appearance, partly combined with a brown crust and at a temperature of 1200 °C the valves became transparent and glassy showing plastic deformation which resulted in the loss of ornamentation details. Fossil conchostracan valves exhibit a similar alteration path allowing a very broad estimate of the thermal effect of the basalt flows on intercalated sedimentary rocks. However, there is a considerable overlap of valve colour, i.e. valves on the same bedding plane are generally assigned to at least two geothermometry classes. Additionally, Tasch (1982) did report some mineralogical data of the enclosing sedimentary rocks but detailed analyses such as illite crystallinity are lacking and only very limited information on palynomorph preservation are available. Samples which are assigned to the two low temperature classes yielded excellently preserved palynomorphs of “unaltered color”. Thus, the valve colour classes are not calibrated against other more reliable maturity parameters. Colour change and the degree of deformation of conchostracan valves may have some potential in thermal alteration studies (Supplementary material 3) but this method is far from being well explored. Consequently, very few studies refer to Tasch's valve geothermometry (e.g. Lana and Carvalho, 2002; Tassi et al., 2013). Reflectance measurements of conchostracan valves in polished blocks or colour determination of dispersed valve fragments in palynological slides using transmitted light may be more promising than the broad description of valve changes as proposed by Tasch (1982). In general, conchostracans are quite rare in the geological record compared to other zooclasts used in thermal maturity studies but in restricted environments they can occur in large numbers typically making up monospecific assemblages which are preserved on bedding planes or in thin layers. McDonald (2007, pp. 160–162) reported numerous conchostracan fragments concentrated in Middle Devonian bituminous

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laminites and associated with filamentous and prasinophyte alginites and sporinite. However, the general scarcity of conchostracans, lack of a universally accepted colour scale as well as deformation classification and lack of calibration against standard maturity parameters limit this method to curiosity-driven research projects only. The carapace valves of extant Spinicaudata, one of the suborders of the Diplostraca and belonging to the conchostracans, are weakly mineralized and mainly composed of calcium phosphate and a varying amount of chitin. In exquisitely preserved Maastrichtian spinicaudatans Stigall and Hartman (2008) found C contributing a mass fraction of 10–24%. We are not aware of any study investigating chemical modifications of conchostracan valves during thermal maturation. 2.7. Ichthyoliths 2.7.1. General remarks Ichthyoliths, i.e. skeletal, disarticulated remains of fish, are made of calcium phosphate and may include an organic component. Ichthyoliths are generally smaller than 2 mm and comprise teeth, fin spines, dermal denticles and scales. The earliest undisputed ichthyoliths are Mid Ordovician in age, however, they may extend into the late Cambrian and range into the Holocene. They are found in marine and nonmarine successions and may be very abundant in lag deposits. Ichthyoliths are robust particles and are therefore concentrated due to winnowing or during periods of low accumulation rates. However, they are very rarely recorded during routine organic petrological studies (Creaney, 1980; McDonald, 2007). Ichthyoliths are useful in biostratigraphy and biofacies analyses (e.g. Johns et al., 2006). As they are composed of biogenic apatite, the primary oxygen isotope signal may be preserved provided that diagenetic recrystallization did not alter the original isotope composition. In particular the dense hypermineralized enameloid is considered to record the primary oxygen isotope signal while the basal tissue is more susceptible to recrystallization because of numerous pores and canals (Barham et al., 2012; Fischer et al., 2013). 2.7.2. Colour alteration and mineralogical changes of ichthyoliths during diagenesis Ichthyoliths from immature sedimentary rocks are yellow and become progressively darker with increasing thermal maturity. At very high rank, e.g. adjacent to sills, they are opaque white and finally clear and brittle. Hence, they show a strikingly similar colour path during maturation as is seen in conodonts. This colour change is regarded to be irreversible (Tway, 1982; Tway et al., 1986). In open-air heating experiments, following the conditions applied by Epstein et al. (1977) for conodonts, Tway et al. (1986) heated three ichthyolith groups (palaeoniscoid teeth and scales, placoid scales) as well as conodonts and recorded the changes in colour (for experimental details see Tway, 1982). While the variability in colour of the conodonts was low, that in ichthyoliths was disappointingly high. Although there was a general trend towards darker and more reddish hues, the high variability prevents precise colour classification of the artificially heated ichthyoliths. The considerable colour variability within a single sample is attributed to the porosity in ichthyoliths which leads to combustion of organic matter. Conodonts and the tips of some ichthyoliths employed in the experiment, by contrast, are very dense and display only a slight colour variation. Despite the fact that the heating experiments by Tway et al. (1986) raised doubts about the practicability of ichthyoliths as maturity parameter, several studies (Supplementary material 3) successfully applied these microfossils to decipher complex tectonic structures, basin evolution (Johns et al., 2006, 2012), thermal anomalies (Johns et al., 2000, 2012) and regional maturation patterns for hydrocarbon exploration (Johns et al., 1999a, 1999b, 2000). In these publications the ichthyolith alteration index (IAI) was introduced which closely follows CAI values. Maturity data deduced from ichthyolith colour were partly underpinned by CAI and foraminifera alteration analyses and in combination with a high-resolution stratigraphy offered a reliable basis for the

reinterpretation of the deformation style at a convergent accretionary margin (Johns et al., 2012). Obviously, it is essential to use ichthyoliths with simple form, dense dentin (preferably enamel) and of the same size and shape, and preferably lacking canals as well as pores which make ichthyoliths vulnerable to permineralization (Barham et al., 2012), resulting in a non-thermal induced discolouration. These criteria exclude the use of bone as an indicator of thermal maturity. Bones are highly porous, turn black in colour and simultaneously loose histological details when exposed to a heat source (e.g. dykes; Janvier et al., 2007). However, hydrocarbons tend to stain the porous tissue which results in a colour change irrespective of rank. Under ultra-violet light, bone fragments with trapped hydrocarbons show a bright yellow fluorescence and oil expulsions during excitation (Creaney, 1980). Another drawback is that in comparison to conodonts, ichthyoliths are more rapidly attacked by fungi and bacteria which affects colour (Johns et al., 1999a; Koot et al., 2013; Tway et al., 1986). This may explain why in a single sample ichthyoliths of different colour are frequently present. In this case the reworking of thermally more strained ichthyoliths from older deposits must be considered as phosphatic microfossils are physically and chemically more robust than e.g. aragonite and calcite tests. Distinguishing reworked from indigenous ichthyoliths is of utmost importance when these fossils are used in a biostratigraphic framework. Besides the application of ichthyoliths in thermal maturity studies (Supplementary material 3), their colour is a useful indication whether oxygen and strontium isotope composition is unaffected by diagenesis and can still serve as an archive for palaeotemperature, salinity and for strontium isotope chronostratigraphy (Johns et al., 2012; Žigaitė et al., 2010). So far in ichthyoliths only the colour change has been used as a thermal maturity parameter. Creaney (1980) reports a low reflectance of 0.2% in vertebrate bones, most probably isolated from immature to marginally mature shales and a very variable fluorescence which ranged from no fluorescence to dull brown and yellow–green. Bright fluorescence was associated with absorbed hydrocarbons. Fluorescence colour and intensity of fossil bones is possibly not only dependent on maturity but also on taphonomic factors and fossilization processes (e.g. Eichler and Werneburg, 2010). In some sediments, ichthyoliths and bones are quite rare so that the application of these faunal remains in routine maturity studies is hindered. However, this drawback can be overcome when ichthyoliths and bones are concentrated using standard palaeontological techniques, embedded in epoxy resin and polished for reflectance measurements.

2.7.3. Chemical composition of organic matter in ichthyoliths Knowledge about the organic preservation of ichthyoliths is extremely limited. Using Curie Point-pyrolysis–gas chromatography/ mass spectrometry and TMAH assisted thermochemolysis–gas chromatography/mass spectrometry, Gupta et al. (2008) investigated a well preserved fish scale and a fish bone from the Barremian Las Hoyas site, a freshwater Konservat-Lagerstätte (fossil-lagerstätte with exceptionally well preserved organisms often including soft tissue preservation). The organic component in both specimens is similar and revealed an aliphatic polymer. Collagen-derived biomarkers such as diketodipyrrole are lacking. The long chain acyl moieties range from C8 to C22 with a maximum abundance of C16 and C18 and an even over odd predominance. As the distribution of fatty acid methyl esters in fossil ichthyoliths is very similar to that of modern fish scales, the aliphatic moieties in the fossils are probably the result of in situ polymerization of the original labile aliphatic compounds. A postmortem migration of the aliphatic component from the surrounding sediment into the ichthyoliths is excluded by Gupta et al. (2008) but following e.g. de Leeuw et al. (2006, p. 209) “migration of aliphatic moieties into, and their condensation onto the macromolecule might occur, e.g. by oxidative polymerization”. A systematic comparison of the geomacromolecular composition in ichthyoliths of various thermal maturities is lacking.


2.8. Structureless and structured kerogen In palynological preparations structureless kerogen includes among others amorphous organic matter (AOM) and structured kerogen comprises phytoclasts (mainly fungal and plant debris), palynomorphs (e.g. organic-walled phytoplankton, mio- and megaspores, zoomorphs with chitinozoans and scolecodonts) and zooclasts (arthropod and graptolite debris) (e.g. Tyson, 1995, Fig. 20.1). It should be noted that the percentage of AOM in palynological slides is usually much greater than that of structureless organic matter in incident light microscopy performed on sections cut perpendicular to bedding. Such unstructured organic matter is most common in upwelling sediments (Littke and Sachsenhofer, 1994), whereas in most other sedimentary environments, wellstructured macerals predominate, including small but distinct fragments such as liptodetrinite, vitrodetrinite (detrovitrinite) and inertodetrinite. In palynological preparations graptolites are rarely encountered and more often observed in polished blocks. Graptolites are therefore addressed in a separate chapter (see Section 2.3). All other constituents are regularly found in palynological slides. Acritarchs, organic-walled dinoflagellate cysts, miospores and chitinozoans are of utmost importance in biostratigraphy and thus play an important role in oil, gas and coal exploration, geological mapping and in academic research. Optical palaeotemperature parameters such as colour estimates and quantitative colour measurements (including fluorescence) do not require any further processing techniques than those used for palynostratigraphy or palynofacies analyses. Thus, these parameters can all be determined using the one set of slides which makes this method very cost-effective. Detailed laboratory protocols for the isolation of kerogen concentrates and palynological preparations have been presented in many

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publications (summarized e.g. in Brown, 2008; Suárez-Ruiz et al., 2012; Traverse, 2008; Vandenbroucke and Largeau, 2007) and are not repeated here. Special procedures are needed when large and brittle palynomorphs (e.g. Proterozoic acritarchs) have to be recovered. This requires less vigorous handling such as the crushing of the rocks to larger fragments as usual (up to 0.5–1 mm), decanting instead of centrifuging and also filtration by hand instead of using semiautomatic equipment (e.g. Grey, 1999). Large and more robust palynomorphs such as chitinozoans and scolecodonts may be rare in some samples and this necessitates the processing of larger samples with further concentration by sieving and hand-picking. Extraction techniques for these large palynomorphs and plant/animal debris follow well established micropalaeontological procedures (e.g. Green, 2001; Pearson and Scott, 1999). The hand-picked specimens can be utilized for chemical analyses or for reflectance measurements after embedding and polishing. Especially when subsequent organic–geochemical studies such as μFTIR spectroscopy or pyrolysis–gas chromatography/mass spectrometry are intended, the processing of palynomorphs is a critical issue and particular precaution must be adopted to avoid any chemical alteration. Standard palynological processing of siliciclastic sedimentary rocks usually involves demineralization by hydrochloric and hydrofluoric acid. The breakdown of the mineral matrix can be accelerated by heating in a sand bath. However, introducing heat during processing may be questionable and usually is not necessary, particularly since even large rock fragments disaggregate after some time provided that the hydrofluoric acid is renewed several times. Instead, to be on the safe side, the beakers can be placed in a water bath in order to carry off the heat released during the exothermic reaction after the sample is poured with hydrofluoric acid. It has been reported that acid digestion does have

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Fig. 13. Spectral distribution of fluorescence intensities and spectral parameters of miospores from the Permian of the Southern Alps. (a) Comparison of miospore fluorescence spectra in polished block and after palynological processing; (b) Comparison of miospore fluorescence spectra from the Bolzano Volcanic Complex and the overlying Gröden Sandstone. Miospores from the Gröden Sandstone show higher thermal maturity.

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an effect on the fluorescence colour of miospores due to the generation of heat during reactions (e.g. McPhilemy, 1988). Compared to polished blocks, a red shift has been observed for alginite in kerogen concentrates which was attributed to the still rather mild acid treatment (Mendonça Filho et al., 2010). No distinct difference in fluorescence properties has been noted between miospores in whole rock and after palynological processing (Hartkopf-Fröder et al., 2001) but this comparison was based on only one sample (Fig. 13a). Short-term heating is believed not to alter miospore colour (Marshall and Yule, 1999). Hydrochloric/hydrofluoric acid maceration probably does not alter the chemical composition except if the material is thermally very immature and still contains polysaccharide moieties (van Bergen, 1999). For unconsolidated sediments a non-acid preparation technique which uses sodium hexametaphosphate [(NaPO3)6] has been developed (Riding and Kyffin-Hughes, 2011). This method, however, includes boiling the sample in an aqueous sodium hexametaphosphate solution and is not effective for indurated sedimentary rocks. The additional treatment with hydrogen peroxide must be excluded (step 5 in Riding and Kyffin-Hughes, 2011). It is still unclear whether or to what extent this procedure alters the chemical composition of palynomorphs. Acetolysis, alkali treatment and any oxidizing reagents such as nitric acid and Schulze's solution (70% nitric acid supersaturated with potassium chlorate) must be strictly avoided regardless of whether the samples are processed for organic-geochemical or for colour analyses, as these treatments drastically change the chemical composition and colour of palynomorphs (e.g. van Bergen, 1999). For this reason, the extensive studies on colour alteration of AOM, acritarchs, dinoflagellates, miospores, chitinozoans and scolecodonts by Correia (1967, 1969, 1971) are dubious as he stipulated the use of nitric acid for standard extraction. Also boiling the residue in potassium hydroxide as recommended by Gutjahr (1966) can result in distorted palynomorph colours (Staplin, 1969). The exclusion of oxidants and alkali treatment implies that colour analyses cannot be performed on miospores macerated from coals which can only be disintegrated by applying these chemicals. For reflected light microscopy, studies on whole rock sections cut perpendicular to bedding are usually preferred, because all organic particles can be examined in the same orientation and in their natural position within the rock. However, at low levels of organic matter content, studies of kerogen, palynomorph and zooclast concentrates are more efficient. Sieving with appropriate mesh, if needed complemented by ultrasound to remove AOM, will greatly increase the microfossil yield (e.g. Marshall, 1995). These concentrates are either embedded in resin blocks and polished which is the standard for coal petrographic analysis (e.g. Suárez-Ruiz et al., 2012; Taylor et al., 1998) or the concentrates can be mounted in polished thin sections (Hillier and Marshall, 1988). Embedding small objects such as chitinozoans or scolecodonts in resin blocks can be challenging especially when they are rare, friable or have to be oriented in the block. For this problematic material resins with short pot lives should be employed. A two-stage embedding system in which the organic material is first stuck to the mould base by a thin film of e.g. cyanoacrylate and then covered by epoxy resin prevents floating of the specimens and fixes them on a single plane surface (Jones and Rowe, 1999). Compared with the traditional polished block method, the one described by Hillier and Marshall (1988) requires smaller amounts of organic matter, is more rapid and enables observation in reflected and transmitted light microscopy. Under incident light, structureless organic material and liptinite macerals can best be observed using a fluorescence mode (Taylor et al., 1998). However, quantifying the area percentage of such unstructured organic matter is difficult, because the relative percentage of submicroscopic organic matter and mineral matter cannot be deduced easily just based on fluorescence colour and intensity. Bitumen impregnation (within the oil window) greatly changes fluorescence colour and intensity, i.e. solvent extracted surfaces of polished sections show much different fluorescence of alginites than untreated surfaces (Littke et al., 1988).

2.8.1. AOM 2.8.1.1. General remarks. AOM is especially abundant in many highly oilprone source rocks deposited under oxygen-deficient bottom water conditions in particular in upwelling areas along ocean margins (Littke and Sachsenhofer, 1994). AOM are organic masses which show no cellular structures and have no sharp outline. It can be derived from strongly degraded tissue, but also from condensation/polymerization of smaller organic molecules (Hedges et al., 2000). AOM in aquatic sediments is probably largely derived from microbially and chemically altered phytoplankton and algae but microbial mats (cyanobacteria, thiobacteria) and faecal pellets most probably also contribute to AOM (Batten, 1996a; Tyson, 1995). Many authors have included AOM in more or less complex classification systems of sedimentary organic matter (see Senftle et al., 1993; Thompson and Dembicki, 1986). In studies performed in reflected light, AOM (also used terms: unstructured organic matter, bituminite, amorphinite) is characterized by low reflectance (at low maturity levels) similar to other liptinite macerals. Bituminite differs from alginite by the lack of a well-defined shape and has reddish to dark brown fluorescence (Taylor et al., 1998; Teichmüller and Ottenjann, 1977). It commonly occurs in form of lenses of irregular shape (Littke et al., 1988) and as matrix bituminite which mergers with the groundmass (Creaney, 1980). This mineral-bituminous groundmass is commonly observed in petroleum source rocks, often showing bright yellow to brownish/reddish fluorescence (e.g. Creaney, 1980; Sachse et al., 2014; Thompson and Dembicki, 1986). Quantification of the organic matter content within this groundmass is, however, difficult, although it can contain the bulk of the total organic matter (see Table 4.9 in Taylor et al., 1998). Due to preparation techniques, the percentage of AOM in palynological transmitted light studies is commonly even larger than in petrographic studies. Not only all the AOM of petrographic studies, but also the bulk of small liptodetrinite and lamalginite falls into the AOM category in transmitted light. Colour and descriptive terms such as granular, spongy, fibrous, membraneous or finely disseminated have been considered (e.g. Batten, 1983, 1996a) but fluorescence microscopy seems to be more promising to categorize amorphous matter as long as it is thermally immature to early mature (Senftle et al., 1987). With increasing degradation and maturation the fluorescence of AOM diminishes. Observation under fluorescent light facilitates to distinguish between strongly fluorescing AOM of aquatic origin from terrestrially derived material. A very useful qualitative scheme for AOM classification and the preservation of immature kerogen has been presented by Tyson (1995, Table 20.2). The 6-point scale is based on the visual assessment of fluorescence intensity. Thompson and Dembicki (1986) presented another easy to use typification which distinguishes four AOM types on the basis of differences observed in transmitted and reflected white light as well as under fluorescent light. Both classifications are designed for easy and rapid assessment of kerogen in source rocks. The kerogen types can be correlated with elemental and organic geochemical parameters and thus offer information on the hydrocarbon generation potential (inert, gas prone, oil-prone, very oil-prone). Senftle et al. (1987) also proposed to distinguish between fluorescing and non-fluorescing AOM. These authors could show that quantity of fluorescing AOM plus liptinite correlates positively with Rock-Eval hydrogen index values indicating a significant petroleum generation potential of this material. A more sophisticated approach is the examination of AOM using scanning electron, transmission electron and atomic force microscopy whereby the nature of ultralaminae in AOM can be deciphered. Pacton et al. (2008, 2009, 2010) clearly demonstrated that in some AOM these ultralaminae do not only originate from algal cell walls but also from bacterial cell walls, thylakoid membranes and filamentous organisms. Hence, detailed characterization of AOM may contribute to a more in-depth palaeoenvironmental analysis as e.g. thylakoids are an


indicator for benthic photosynthesis and thus for palaeobathymetry (Pacton et al., 2008, 2009). 2.8.1.2. Colour alteration and fluorescence of AOM during diagenesis. The omnipresence of oil-prone AOM in petroleum source rocks prompted the idea to apply colour alteration and fluorescence colour in thermal alteration studies, particularly since the measurements are taken on that material which is most interesting in palynofacies analyses addressing exploration issues. Essentially, three concepts have been developed: visually assessed colour estimates, the quantitative Transmittance Color Index (TCI) and fluorescence properties (intensity and colour, both visually and quantitatively assessed). In contrast to methods which use fossils as a maturity parameter, AOM does not offer any biostratigraphic control, i.e. reworking and sample contamination are rarely recognized but may seriously influence the measurements. Peters et al. (1977) heated amorphous kerogen from recent marine sediment at various temperatures and length of time and subsequently determined the colours using transmitted light microscopy and published colour standards. The initial kerogen colour was yellow and turned brown and finally black upon increasing temperature and time. During these experiments the maximum n-alkane generation, reflecting a maturity level corresponding to the peak of oil generation, was observed in the dark brown to very dark brown colour range. In order to avoid false colour assessments due to varied kerogen thickness, colours were recorded from kerogen of similar size and thickness. Peters et al. (1977) concede that kerogen thickness can be problematic and they themselves rate this method as a “quick and dirty” technique, which however also has its advantages as it is rapid, inexpensive and allows a rough thermal maturity estimate. Visually assessed AOM colour data have been applied in numerous source rock studies. The general practice is to make use of the Pollen/ Spore Color “Standard” of Pearson (1990) in order to objectivize the colour identifications as far as possible. This method is also successful in sediments of Cambrian and Precambrian age or in mineralized successions where sporomorphs are either lacking or have been destroyed due to hydrothermal fluids and high heat flow (e.g. Christiansen et al., 1989; Parnell and Janaway, 1990; Robbins, 1983; Robbins et al., 1990). Marshall et al. (1985) ascertained the heating effect of a dyke on the surrounding sedimentary rocks by using AOM colour change. Obviously, the heat from the dykes is negligible compared to burial maturation. A quantitative approach to analyse AOM colour was presented by van Gijzel et al. (1992) and Robison et al. (2000). They defined the new parameter TCI as the average wavelength value in nanometers of the transmittance spectra of AOM particles. In an earlier publication (van Gijzel, 1990) this parameter was also referred to as TSI (Transmittance Spectral Index). The TCI values are approximately 570 nm for very immature AOM and increase up to 640–670 nm for postmature AOM (all measurements in transmitted tungsten light; Robison et al., 2000). As in visual AOM colour determination, the thickness of AOM particles influences the TCI resulting in some scattering of the values. Hence, care must be taken to analyse particles of similar thickness. The method proved most successful in oil-prone source rocks and at immature to postmature hydrocarbon stages. While the objective spectral TCI measurements overcome the drawbacks of visual colour assessments, problems caused by reworking, contamination and impregnation from mobile bitumen cannot be precluded. Nevertheless, a good linear correlation between TCI and vitrinite reflectance has been demonstrated enabling the conversion of TCI values into equivalent vitrinite reflectance values. The same is true for the TCI and TAI correlation (Robison et al., 2000). The main advantages of the TCI is that it can be applied to thermal maturity analyses of sedimentary rocks in which vitrinite and sporomorphs are rare or lacking (e.g. source rocks and lower Palaeozoic and Precambrian sedimentary rocks). Additionally, only slides with unsieved kerogen concentrates need to be prepared. Visually determined fluorescence colours are frequently applied in source rock studies and to characterize kerogen preservation and

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quality with regard to hydrocarbon generation. Based on a large dataset of AOM fluorescence properties and HI Tyson (2006) demonstrated the relationship between both parameters. Fluorescence data can best be presented in a ternary diagram to portray the proportions of the different kerogen fractions (Mustafa and Tyson, 2002). 2.8.1.3. Chemical alteration of AOM during diagenesis. Diagenetic alteration of AOM is a very complex process depending on many different parameters. However, to summarize the chemical characterization, generation and alteration of AOM in detail goes beyond the scope of this review (for an overview see e.g. Hedges et al., 2000; Rullkötter and Michaelis, 1990; Vandenbroucke and Largeau, 2007). Kinetic studies on petroleum generation characteristics have been performed for many source rocks in which AOM is present in great quantity revealing highly variable petroleum generation characteristics. These kerogens are also highly variable in terms of Rock-Eval hydrogen index values, but usually fall into the fields of type I, type II or type II/III kerogen (see Table 4.9 in Taylor et al., 1998; HI between 200 and 900). Some of them such as the lacustrine Nördlinger Ries Shale (Germany) or the Cretaceous marls of the Tarfaya Basin, Morocco, are extremely rich in organic sulphur, hinting at an early diagenetic restructuring of kerogen by vulcanization processes, whereas others have low organic sulphur contents such as the Messel Shale and Posidonia Shale (both Germany) or Pennsylvanian black shales (USA). Thus the chemical and structural variability of AOM is probably enormous, covering almost the entire range that we know from better structured kerogen types. 2.8.2. Acritarchs Acritarchs are fossilized, organic-walled cysts of unicellular protists that cannot be assigned to known groups of organisms (Strother, 1996). The acritarchs comprise a variety of organic remains such as resting cysts of marine phytoplankton and arthropod eggs. Originally, the term was introduced as an informal taxonomic, heterogeneous and polyphyletic category to accommodate cysts and cyst-like organic microfossils remaining after most of the hystrichospheres were identified as Dinophyceae (Evitt, 1963; Martin, 1993; Playford, 2003; Traverse, 2008). Most acritarchs consist of a central body modified by surficial sculptural elements. The inception of acritarchs is dated as Mesoarchean (Javaux et al., 2010). They are stratigraphically important particularly for the Palaeozoic but are also present in modern sediments. Acritarchs have been recorded from brackish water and freshwater sedimentary rocks but are most diverse in marine environments where they may be extremely abundant (e.g. up to 105 specimens/g sedimentary rock, Downie, 1958). Detailed reviews on acritarch morphology, palaeoecology and biostratigraphy have been published by Martin (1993), Moczydłowska et al. (2011) and Playford (2003). 2.8.2.1. Colour alteration and fluorescence of acritarchs during diagenesis. Acritarchs display colour changes with increasing thermal maturity which is best observed in thin, smooth-walled forms without inner body such as Veryhachium. Thermally unaltered acritarchs are colourless and transparent and pass through yellow, brown and ultimately black opaque with progressing temperature (e.g. Dorning, 1986; Duggan and Clayton, 2008; Martin, 1993). Hence, the colour maturity path of acritarchs is similar to that of miospores (see below) but not identical (Figs. 14, 15). Acritarchs display paler and weaker colours when compared to miospores and black opaque conditions are attained at higher temperatures than in miospores (e.g. Collins, 1990; Duggan and Clayton, 2008; Hartkopf-Fröder et al., 2004; Playford, 2003). As acritarchs occur in vast amounts in marine and in pre-Devonian sedimentary rocks where vitrinite is rare or absent, they can be of particular value in maturation studies. Surprisingly, they have rarely been used to obtain an approximation of the thermal maturity and in this respect their potential is far from explored. Based on the change in leiosphere colour Legall et al. (1981) proposed a 5-point scale, the Acritarch Colour Alteration Index (AAI). In the oil window leiospheres

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Fig. 14. Correlation between miospore and acritarch colour (both determined by using the SCI scale) superposed on hydrocarbon stages. Colour change in acritarchs is more pronounced at higher thermal maturity while miospores have the major colour change in the immature to middle mature stage. Modified from Collins (1990).

show a rapid colour change, i.e. from translucent/light yellow to dark brown. However, at least most of the Leiosphaeridia species are now assigned to fossil phycomata of prasinophycean algae (Guy-Ohlson, 1996; Strother, 1996) and hence may display a different colour maturity path than that of acritarchs. Obviously, leiospheres are destroyed at lower temperatures than acanthomorphitic acritarchs are (Dorning, 1986; Duggan and Clayton, 2008) indicating differences in wall chemistry and that, indeed, Leiosphaeridia species used by Legall et al. (1981) may rather be related to prasinophycean algae than to acritarchs. A

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Spore colour index [SCI]

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different, 4-point AAI scale was used by Williams et al. (1998) who also correlated TAI, fluorescence as well as vitrinite and graptolite reflectance data. Collins (1990) applied the 10-point Spore Colour Index (SCI) of Fisher et al. (1980) to characterize both acritarch and miospore colours. The most thorough study on acritarch colour alteration was performed by Duggan and Clayton (2008) who published a diagram of acritarch colour change based exclusively on Veryhachium and using RGB intensities. The advantage of this approach is that exclusively smooth, thin-walled acritarchs of uniform wall thickness are selected, reducing colour variation due to structural or sculptural differences. However, RGB intensities are influenced by e.g. microscope and camera settings impeding interlaboratory comparability. Another attempt to use a quantitative technique (μ-spectrophotometry and presentation of the complete spectrum in the international colour system of the Commission Internationale de l'Éclairage) is that of HartkopfFröder et al. (2004) which, however, is limited to a very narrow maturity level at approximately 0.5–0.6% VRr (Fig. 15). Similar to miospores (see below) the colour change in acritarchs is paralleled by a reduction in size of the vesicles and a thinning of the processes (Di Milia, 1991) with increasing thermal alteration. Acritarchs still show some fluorescence at a higher rank than miospores, e.g. the fluorescence of Veryhachium, a common, thinwalled genus with simple morphology, ceases between 1.5 and 2.0% VRr while in miospores it extinguishes at approx. 1.35% VRr (Duggan and Clayton, 2008). Hence, qualitative fluorescence of acritarchs can be routinely used as a maturity index in the immature to postmature hydrocarbon stages. Fluorescence μ-spectrophotometry which is also a quantitative method of analysing the relative fluorescence in the 400–700 nm range, was applied by Obermajer et al. (1999a). They proved that acritarch fluorescence is dependent on thermal evolution and that the spectra show a progressive shift from yellow near the oil birth line to red near the peak oil generation. Occasionally, acritarch reflectance has been studied as a parameter for thermal maturity. However, this approach never achieved importance in organic petrology (Goodarzi et al., 1985) probably because most Palaeozoic to Cenozoic acritarchs are small and have very thin vesicle walls. However, Precambrian acritarchs reach up to approx. 200 μm. Chuaria, interpreted as a thick-walled sphaeromorph acritarch or a multicellular eukaryotic alga (Xin et al., 2011), has a maximum size of up to some millimetres.

40

acritarchs (Veryhachium spp.)

20 20

CIELAB a*

CIELAB chroma

0 30

CIELAB chroma

CIELAB a* 0

0 0

miospores (Diducites mucronatus and D. poljessicus)

20

0

30

60

0 0

30

0

30

60

Fig. 15. CIELAB chromaticity and colour tone diagrams of acritarchs and miospores. Acritarchs display weaker colours than miospores. Borehole Refrath 1, Rhenish Massif, Germany. Modified from Hartkopf-Fröder et al. (2004).


Reflectance of these large sized organic fossils may serve as a thermal maturity parameter in Precambrian sedimentary rocks. Apart from the above-mentioned publications, acritarch colour and fluorescence were used in various thermal maturity studies (e.g. Al-Ameri, 2010; Belka, 1993b; Moczydłowska, 1988; Szczepanik, 1997). Especially interesting is the observation by Moczydłowska and Vidal (1992) who found acritarchs in phosphorites showing darker colours than in accompanying shales and mudstones. The authors relate these colour discrepancies to irradiation from natural radioactive decay in the phosphorites. A similar phenomenon has been reported by Stasiuk (1994a). Alginite from the Bakken Formation (DevonianMississippian, Canada) shows a red-shift in fluorescence and hence indicates a higher thermal maturity when in close association with phosphatic particles enriched in radioactive elements. 2.8.2.2. Chemical composition of acritarchs. The macromolecular composition of acritarchs has not yet been explored in detail except for numerous studies dealing with Precambrian acritarchs to unravel their biological affinity. μ-Raman spectroscopy, μ-FTIR spectroscopy, synchrotron radiationbased μ-FTIR and pyrolysis–gas chromatography/mass spectrometry have been used (e.g. Carter et al., 2007; Dhamelincourt et al., 2010; Javaux and Marshall, 2006; Kempe, 2003; Marshall et al., 2005; Olcott Marshall and Marshall, 2015; Schiffbauer et al., 2007). Talyzina et al. (2000) examined Early Cambrian acritarchs applying gas chromatography/mass spectrometry/mass spectrometry (GC/MS/MS) and metastable reaction monitoring–gas chromatography/mass spectrometry (MRM–GC/ MS) and based on the presence of dinoflagellate-related biomarkers (dinosteranes and 4α-methyl-24-ethylcholestane) they suggested a dinoflagellate affinity for at least some acritarch genera. Following a μ-FTIR study, an unusual large acritarch species from the Givetian of Libya is composed of an algaenan similar to that from Botryococcus braunii (Steemans et al., 2009a). The fate of post-Precambrian acritarch biopolymers altered due to thermal maturation has never been elucidated. 2.8.3. Organic-walled dinoflagellates 2.8.3.1. General remarks. Originally, microfossils now classified as fossil dinoflagellate cysts were included in the hystrichospheres until it was demonstrated that some of the hystrichospheres have excystment openings and show a tabulation pattern similar to modern dinoflagellates (Evitt, 1963; for a review see Fensome et al., 1996). Dinoflagellates are an important part of the phytoplankton and as primary producers they may be present in vast amounts, i.e. one gramme may yield up to 5 × 106 specimens (Traverse, 2008). Along with acritarchs and prasinophytes they constitute the majority of the fossil organic-walled marine microflora. The resistant organic-walled dinoflagellate cysts (single-celled protists) occur primarily in marine strata of Mid-Triassic to Holocene age, although few fresh water fossil forms are known. The greatest diversity and abundance was reached during the Jurassic and Cretaceous. There is still some debate whether dinoflagellates have a pre-Mesozoic record. Two Silurian and Devonian genera, Arpylorus and Palaeodinophysis, are morphologically similar to dinoflagellate cysts but assignment of these genera to dinoflagellates is controversial (Fensome et al., 1996, 1999) and was recently rejected for the Silurian genus Arpylorus (for a summary see Medlin and Fensome, 2013). Based on morphological, ultrastructural and biogeochemical evidence the earliest putative dinoflagellate cysts have been reported from Mesoproterozoic rocks of China (Meng et al., 2005). Interestingly, triaromatic dinosteroids, which are derived from dinosterols, have been detected in extracts of Precambrian rocks. Dinosterols are generally believed to be biogeochemical markers for dinoflagellates. Therefore, it was concluded that dinoflagellates or closely related precursors which were capable of biosynthesizing these specific compounds were thriving in the oceans long before the first undisputed organic-walled and resistant dinoflagellate cysts appeared in the fossil record (Li et al., 2012;

97

Moldowan et al., 1996). This assumption was further reinforced by the discovery of dinosteranes in acritarch concentrates of early Cambrian age (Talyzina et al., 2000). In fact, molecular phylogenetic data give evidence for a dinoflagellate origin in the latest Proterozoic, but either these pre-Mesozoic representatives did not produce fossilizable cysts or their cysts still remain unrecognized and classified as acritarchs (Medlin and Fensome, 2013). As not all organic-walled dinoflagellate cysts are resistant enough to be preserved and pre-Mesozoic forms may have left no or only meagre fossilized evidence of their existence, the palynological record of dinoflagellates seems to be rather incomplete. 2.8.3.2. Colour alteration and fluorescence of organic-walled dinoflagellates during diagenesis. It is claimed that organic-walled dinoflagellates show the same colour maturity path as acritarchs (Dorning, 1986). However, optical changes of dinoflagellates during maturation are poorly documented and have never properly been used in maturation studies with the exception of an unpublished Ph.D. thesis (Milton, 1993). The thesis presents a vast amount of data produced by μ-spectrophotometry and presented in CIELAB diagrams. However, in most plots dinoflagellates were not separated from miospores and as these two groups show pronounced differences in colour during maturation, the interpretation of the majority of Milton's data is difficult. Plots which solely depict dinoflagellates demonstrate that wall thickness does have an influence on colour i.e. thick-walled specimens are darker and thus seemingly more mature than thin-walled dinoflagellates. Two- or multi-layered cyst walls often show differences in colour of the wall layers. In general, unaltered dinoflagellate cysts are colourless to pale yellow and change colour with increasing maturity through brown, grey and finally black (Correia, 1967; Dorning, 1986). Compared to miospores they display paler colours. Using the colour change of dinoflagellate cysts as a thermal maturity parameter is further hampered by pronounced colour differences in unaltered cysts and selective preservation. Some cysts, most of them belonging to the genera Islandinium and Echinidinium, have brown coloured walls even when unaltered and well preserved. These cysts have been informally described as “round brown spiny cysts” (Radi et al., 2013) and this group should be excluded in maturity and provenance studies (e.g. when in situ and reworked material has to be distinguished). Selective preservation is rather well studied in fossil dinoflagellates (for a summary see e.g. Ellegaard et al., 2013). Some dinoflagellate cysts are extremely resistant while others are highly vulnerable to degradation or are not fossilizable. Oxygen concentration in bottom water and sediment is crucial for the preservation of cysts with no or nearly absent degradation under anoxic conditions and high degradation rates in oxic environments. Even physical (e.g. sonification, excessive heating) and chemical treatments (e.g. acetolysis, oxidation, KOH) occasionally applied in palynological processing may selectively damage or destroy dinoflagellate cysts, resulting in the enrichment of more resistant specimens such as reworked cysts which have undergone higher thermal maturation (e.g. Mertens et al., 2009; Schrank, 1988). Partially degraded cysts or cysts damaged because of inappropriate processing techniques are corroded, have thinner walls and may be torn or heavily folded which makes them unsuitable for colour estimates. Hence, care must be taken to use only well preserved cysts with a onelayered wall preferably of the same genus and not to apply too harsh processing techniques. Only few studies have used dinoflagellate cyst colour as the only rank parameter. Especially in source rock studies, colour estimated in transmitted light has been combined with fluorescence data of lipid-rich organic matter (see below). However, based on cyst colouration, maturity assessments can easily be carried out provided that the caveats outlined above are properly addressed. No specific colour scale has been proposed for dinoflagellate cysts. As a substitute well introduced scales such as the Pollen/Spore Color “Standard” of Pearson (1990) or the Spore Colour Index (SCI) of Fisher et al. (1980) and Collins (1990) can be consulted to visually determine dinoflagellate colour (e.g. Blažeković Smojić et al., 2009; Firth, 1993; Honigstein et al., 1989).

98


Apart from source rock studies, colouration of organic-walled dinoflagellate cysts has been successfully applied to distinguish reworked specimens from contemporaneous ones and to identify provenance areas as well as drainage routes (e.g. Smelror, 1999). Fluorescence properties of dinoflagellates as a function of thermal maturity have been studied in several publications, mostly in addition to other liptinite macerals such as sporinites, alginites (prasinophycean algae) and cutinites (e.g. Gentzis et al., 1993; Goodarzi et al., 1989). Almost all of these studies ignore the pronounced differences in fluorescence characteristics between photosynthetic and heterotrophic organic-walled dinoflagellate cysts. The overwhelming majority of extant taxa are photosynthetic, mixotrophic or heterotrophic, the latter constituting about one-half of the species (e.g. Protoperidiniaceae, Gymnodiniaceae). Cysts of the non-photosynthetic heterotrophs have pigmented cyst walls, show weak or no visible autofluorescence and are readily destroyed by weak oxidation, while phototrophic species are not pigmented, are highly autofluorescent and are more resistant to oxidizing agents (e.g. Brenner and Biebow, 2001; Feist-Burkhardt, 2009). These intrinsic differences are mirrored by the cyst wall chemistry (Bogus et al., 2014; see below). Whereas the greatly differing fluorescence properties offer an exciting approach to unravel the nutritional mode of fossil organic-walled dinoflagellates, heterotrophic cysts must be excluded when focussing on thermal maturity as weak or no visible autofluorescence of these cysts indicates only the specific trophic mode and not higher thermal stress. Hufnagel (1977), who was one of the first to report on the fluorescence properties of fossil dinoflagellate cysts, only used fluorescence intensities but as no absolute standard existed for this parameter this approach was not followed up. Nevertheless, he demonstrated that with increasing maturity fluorescence intensity decreases. The fluorescence parameters λmax and the red/green quotient Q (relative intensity at 650 nm/relative intensity at 500 nm) shift with increasing vitrinite reflectance to higher values, i.e. from yellow to red colours, but this correlation is only well explored in the immature to peak mature stage (VRr 0.4–0.9%). Compared to the fluorescence of Botryococcus and sporinite organic-walled dinoflagellate cysts show a slower increase of both parameters in this maturity range (Gentzis et al., 1993; Goodarzi et al., 1989). Apart from thermal rank and trophic mode, fluorescence is affected by oxidation and corrosion of the cyst during deposition, early diagenesis, reworking and due to weathering (Waterhouse, 1998). Fluorescence colour estimations and fluorescence quantitative parameters of organic-walled dinoflagellate cysts represent a supplementary tool used to cross check vitrinite reflectance data in thermal maturity studies (e.g. Blažeković Smojić et al., 2009), to discriminate younger in situ cysts from older reworked ones (even multiple episodes of reworking can be detected; e.g. Waterhouse, 1998) and to reconstruct the provenance of glacial sediments (Salzmann et al., 2011). 2.8.3.3. Chemical composition of organic-walled dinoflagellates. Extant and fossil dinoflagellate cyst wall chemistry has recently been reviewed by de Leeuw et al. (2006) and Zonneveld et al. (2008). The resting cysts of modern dinoflagellates are composed of the refractory biopolymer dinosporin. However, dinosporin does not represent a uniform molecule but seems to be composed of several different biopolymers. Only few samples have been studied to date and acid treatment of modern cysts preceding analyses seems to have produced artefacts. It is clear that dinosporin is distinct from sporopollenin and algaenan, indicated e.g. by the absence of alkane/ alkene doublets in pyrolysis analysis as is characteristic for algaenans (Versteegh et al., 2012). The cysts of Lingulodinium polyedrum seem to be composed of ether bond-rich, but CH2/CH3-poor aliphatic polymer rather than cross-linked carbohydrates (Versteegh et al., 2012; Zonneveld et al., 2008). In Scrippsiella ramonii the cyst walls are composed of aromatic and aliphatic moieties but lacking isoprenoid moieties (de Leeuw et al., 2006). A network of linear carbon chains and aromatic rings or carbohydrates make up the biomacromolecule

(Zonneveld et al., 2008). Recent investigations applying μ-FTIR on extant dinoflagellates pointed to differences in the molecular composition between heterotrophic and phototrophic species. For both groups a carbohydrate-based backbone was stated, but dinocysts from heterotrophic species exhibited significant contributions of amide groups lacking in dinocysts from phototrophic species (Bogus et al., 2014). Information on the chemistry of fossil dinoflagellate cysts is also very limited and the macromolecular composition varies from almost aliphatic to almost aromatic substances without any isoprenoids (de Leeuw et al., 2006; Versteegh et al., 2007). The chemical differences of dinosporin have even been revealed by μ-FTIR for three Paleogene morphospecies from the same genus (Apectodinium). The molecular properties ranged from an ether-rich cellulose-like composition to carboxyl-rich moieties (Bogus et al., 2012). However, the biomacromolecule can be intensively altered during early diagenesis, e.g. by invading lipids from the organism itself (Gupta, 2014), by oxidative polymerization (Versteegh et al., 2004) or by early sulphurization (Versteegh et al., 2007), but the influence of temperature, time and pressure on dinoflagellate cyst chemistry during diagenesis is generally highly unexplored to date. 2.8.4. Prasinophycean algae 2.8.4.1. General remarks. Prasinophycean algae are considered to be the most primitive green algae (Guy-Ohlson, 1996). Leiosphaeridia and Tasmanites, ranging from the Precambrian to the Holocene, are two important genera included in the prasinophytes. Prasinophycean algae are cosmopolitan and mainly marine but have also been found in shallow lagoonal and deltaic sedimentary rocks. The tasmanite from Tasmania contains a massive algal bloom of the prasinophycean alga Tasmanites (Guy-Ohlson, 1996). Many black shales with prasinophycean mass occurrences are petroleum source rocks, e.g. the Toarcian Posidonia Shale, although they are rarely the predominant type of organic matter (Littke et al., 1988). 2.8.4.2. Colour alteration and fluorescence of prasinophycean algae during diagenesis. Colour changes of prasinophycean algae have rarely been used in thermal maturity studies, although Stadnichenko (1929) has already observed that when heated in a furnace they begin to darken at 440 °C. Traverse (2008, Fig. 19.1b) published a graph showing the alteration in colour of various kinds of palynomorphs. Algal remains are colourless in transmitted light when unheated but rapidly change colour through brown to black at relatively low thermal maturity (Dorning, 1986). Legall et al. (1981) proposed an AAI which is based on leiospheres, at that time assigned to acritarchs but nowadays at least partly to prasinophycean algae. They noted a pronounced colour modification from light yellow to dark brown in the oil window. At higher temperatures leiospheres become degraded. Although prasinophyte colour modification can be used as a qualitative parameter, this index never became an important tool in thermal maturity studies. In contrast to colour changes of prasinophytes observed in transmitted light the fluorescence properties of algae (= telalginite or lamalginite) are commonly used and well explored to determine the thermal maturity (e.g. Blažeković Smojić et al., 2009; Stasiuk, 1994a). When immature, Tasmanites fluoresces greenish yellow and with increasing maturity the fluorescence colour shifts to bright yellow, red and brown (Alpern, 1987; Araujo et al., 2014; Hackley and Kus, 2015). Within the oil window, bitumen impregnation influences fluorescence colour and intensity (Littke et al., 1988). In addition, irradiation from natural radioactive decay, degradation and weathering of outcrop samples greatly influence fluorescence properties leading to longer wavelength of fluorescence (Littke et al., 1991; Stasiuk, 1994a; Stasiuk and Goodarzi, 1988). Fluorescence of Tasmanites alginites ceases between a stage corresponding to 0.9–1.0% VRr (Taylor et al., 1998) or 1.4% VRr (Araujo et al., 2014; Mukhopadhyay, 1994). The fluorescence change in Leiosphaeridia occurs


99

at earlier stages compared to acritarchs. When using fluorescence μspectrophotometry, λmax and the red/green quotient (Q) can be calculated. With the onset of oil generation a rapid change in both parameters towards higher values is observed, i.e. λmax shifts from 450 nm at 0.5% VRc to 600 nm at 0.8% VRc and Q from 0.5 at 0.5% VRc to 1.0 at 0.8% VRc (VRc calculated from reflectance of chitinozoans). Hence, leiospherids are very sensitive to increasing maturity in the lower half of the oil window (Obermajer et al., 1999a). Compared to other algae and miospores the same is true for Tasmanites, at least in the range of 0.5–0.75% VRr (Littke et al., 1988). However, fluorescence properties of prasinophytes are not only dependent on thermal maturity but vary also between prasinophytes derived from Palaeozoic and Mesozoic sedimentary rocks which may be due to macromolecular differences (Stasiuk, 1994b).

aliphatic building blocks which are connected with ester and ether cross links and therefore these algae show high A- and C-factor values in μFTIR analyses (Dutta et al., 2013). Analysis of the macromolecular composition of prasinophycean algae has been restricted to thermally immature material. Molecular alteration of prasinophycean algae during maturation has never been investigated in detail. However, Schenk et al. (1990) performed pyrolysis experiments on an immature Permian tasmanite concentrate from Tasmania comparing it to immature vitrinite (xylite). The initial hydrogen/ carbon ratio of tasmanite was high (1.6), whereas oxygen/carbon ratio was low (0.1) with most oxygen probably bound in carbonyl or carboxyl groups according to IR spectroscopy. Upon heating, tasmanite lost 85% of its mass (vitrinite 60%), mainly in form of hydrocarbons.

2.8.4.3. Chemical composition of prasinophycean algae. Because of their importance as marine oil-prone source rocks tasmanite and Tasmanites-rich sedimentary rocks have been extensively studied with regard to their molecular composition. Aquino Neto et al. (1992) found high tricyclic terpane concentrations in Tasmanites-rich sedimentary rocks from Tasmania, Alaska, and Brazil. All three samples are immature (estimated maturity 0.3–0.6% VRr). Based on the co-occurrence of Tasmanites blooms and abundant tricyclic terpanes, the authors suggested that these compounds are biomarkers diagnostic for this genus. In fact, Tasmanites-rich oil shales are characterized by tricyclic terpenoids and pure Tasmanites concentrates showed C19–C28 tricyclic terpanes and unsaturated derivatives on laser-pyrolysis–gas chromatography/mass spectrometry analysis (Greenwood et al., 2000). However, not all analytical studies supported the precursor (Tasmanites)-biomarker (tricyclic terpanes) relationship (e.g. Talyzina et al., 2000). Using Curie Point-pyrolysis–gas chromatography/mass spectrometry applied to hand-picked pure concentrates of prasinophytes (thermal maturity Tmax 430 °C, equivalent to 0.6% VRr) Dutta et al. (2006) detected tricyclic terpenoids for Leiosphaeridia but not for Tasmanites. A prominence of n-aliphatic pyrolysis products is typical for both genera (de Leeuw et al., 2006; Dutta et al., 2006). Using μ-IR spectroscopy Kjellström (1968) and Dutta et al. (2013) analysed Palaeozoic Tasmanites and Leiosphaerida. Prasinophytes are composed of

2.8.5. Sporomorphs 2.8.5.1. General remarks. For simplification we use sporomorphs in the sense of Tyson (1995), i.e. they include cryptospores, miospores and megaspores. The sporomorph walls are sporopolleninous and highly resistant to most chemicals, temperature and pressure. Cryptospores are the propagules of the earliest embryophytes. At present, the earliest uncontroversial cryptospores have been reported from the Dapingian (early Mid Ordovician) of Argentina (Rubinstein et al., 2010). Cryptospores have no trilete or monolete tetrad mark and occur as hilate monads, dyads or permanent tetrads. Cryptospores were produced by cryptophytes, minute terrestrial plants belonging to embryophyte to tracheophyte stem-groups and relatives of bryophytes (Edwards et al., 2014). Up to the early Silurian (Llandovery) cryptospores dominated the sporomorph associations. They were then superseded by trilete miospores (Wellman, 2014). Miospores encompass iso- and microspores, prepollen, pollen and functional “megaspores” smaller than 200 μm. The earliest record of trilete miospores is from the Katian– Hirnantian (Late Ordovician) (Steemans et al., 2009b). Miospores can be extremely abundant in sedimentary rocks (up to 4 × 106 pollen grains per gramme). Megaspores are produced by heterosporous embryophytes and by definition are larger than 200 μm. Miospores and megaspores can

Fig. 16. Bisaccate pollen from the Lower Jurassic, Hils Syncline, Germany. (a) borehole Wenzen; (b) borehole Wickensen; (c) borehole Dielmissen; (d) borehole Dohnsen; (e) borehole Harderode; (f) borehole Haddessen. Colour alteration with increasing thermal maturity clearly visible. Reduction in size at highest thermal maturity. For more information on this borehole programme in the Hils Syncline see al Sandouk-Lincke et al. (2013).

100


1000

2000

brownish yellow

brown

black

4000

yellow

Depth [m]

3000

1.0

2.0

3.0

4.0

5000

6000 0.0

5.0

Thermal Alteration Index

Fig. 17. Comparison of TAI data from two different laboratories shows pronounced differences in visual colour assessment. Modified from Brosse and Huc (1986).

survive temperatures and pressures of the blueschist (e.g. Bernard et al., 2007) and even the amphibolite metamorphic facies and have been recovered from carbon-rich paragneisses (e.g. Pflug and Prössl, 1991; Pflug and Reitz, 1992).

2.8.5.2. Colour alteration and fluorescence of sporomorphs. As early as the late twenties and early thirties of the last century, it has been noticed that prasinophycean algae and spores and pollen, for example, change colour when heated (Kirchheimer, 1932, 1933; Stadnichenko, 1929;

Stadnichenko and White, 1926; Zetzsche and Kälin, 1932). Based on the colour alteration Kirchheimer (1934) proposed to use spores and pollen as a “thermometer for coalification”. In those days the most thorough study was that of Kirchheimer (1933) who performed comprehensive heating experiments with Lycopodium spores. He noted a size reduction of these spores due to increasing temperature and the colour alteration he ascertained by using colour charts. Shrinkage of spores during heating and thermal maturation has been observed frequently (Fig. 16) and may be used as an indicator of thermal alteration (e.g. Pierart, 1978). Either way, the size reduction must be taken into consideration in case the diameter of a sporomorph species is indicative of biostratigraphical correlations (e.g. the latest Devonian marker species Retispora lepidophyta) between areas of low and high thermal alteration. In transmitted light the unaltered miospore walls are colourless to light yellow and very transparent but turn dark yellow, orange, brown and black with increasing maturity. It is therefore not surprising that the characterization of miospore colour change – as an easy to apply, fast and inexpensive but somewhat subjective method (Fig. 17) – has been widely used in evaluating the potential of hydrocarbon source rocks and their burial history (Marshall and Yule, 1999). Miospore colour alteration is very well explored and visual estimates have been applied in dozens of studies focussing on thermal alteration, regional tectonics and reworking of miospores for provenance studies. These publications need not be reviewed here in detail. A number of miospore colour scales have been developed (Table 2) but none of these scales has been accepted as standard specification. The first colour scales were based on a 5 or 6 points gradation (Correia, 1967; Staplin, 1969) but later the scales were further refined. Three scales, TAI (Staplin, 1969), SCI (Collins, 1990; Fisher et al., 1980; Haseldonckx, 1979) and SCS (Pearson, 1990), became most popular. Originally, the TAI scale was based on “organic debris, especially plant spores, nonwoody cuticle and amorphous sapropelic debris” (Staplin, 1969, p. 57) but was later mainly used for determination of sporomorph colour. The SCI scale is especially important as it is the most detailed one and is well established in the hydrocarbon exploration industry. In addition, a Karweil type diagram (Cooper, 1977) is available which allows calculation of maximum burial temperature (Fig. 18). An approximate correlation between these three scales and vitrinite reflectance has been published by Marshall (1990a) and Marshall and Yule (1999) and a SCI vs. TAI plot is shown in Fig. 19. In order to overcome subjectivity, several methods have been developed to quantitatively and reproducibly measure the colour of miospores. These approaches include colour image analysis (e.g. Yule et al., 1998) or

Table 2 Visual and quantitative miospore colour scales. * Pearson (1990) uses the abbreviation TAI in his colour chart but to distinguish the Pearson scale from Staplin (1969) the abbreviation SCS (Spore Color Standard) is preferable. Also a numbering of 1–11 is more convenient than Pearsons's original subdivision using + and – (e.g. Hillier and Marshall, 1992; Marshall and Yule, 1999). Spore colour scale

Short term

Gradation of scale

Author

Thermal Alteration Index Pollen/Spore Color “Standard” Spore Colour Index

TAI SCS* SCI

5 points 11 points 10 points with intermediates

Thermal Alteration Scale Thermal Alteration Scale Index of Alteration of Microfossils Colour; Colour Index of Microphytofossils Chevron's Thermal Alteration Index

TAS NEWTAI AMC or ICM

7 points with intermediates 10 points 7 points

Staplin (1969) Pearson (1990) Collins (1990); Fisher et al. (1980); Haseldonckx (1979) Batten (1980, 1996b) Løseth et al. (1992) Rovnina (1981, 1984)

Chevron's TAI E.C.

4 points with numerous intermediates 6

État de Conservation CIE colour space Spore Transmittance RGB colour space Integrated Color Analysis via Spectral Power Distribution Statistical Thermal Alteration Index Palynomorph Darkness Index Absorption Translucency Transmittance at 546 nm

% St ICA/SPD stTAI PDI

Jones and Edison (1979) Correia (1967) Marshall (1991) Yule et al. (1999) van de Laar and David (1998), Yule et al. (1998) Jansonius and Schwab in Batten (1996b) Ujiié (2001) Goodhue and Clayton (2010) Gutjahr (1966) Grayson (1975) Lo (1988)


101

10

9

160

3.

7

140 130 120

6

110

0

8

2.5

7

5

100

6

90

SCI

Maximum temperature [°C]

9

8

150

4

80

5 ?

70

3

60

4

2.0

50

3

Smith (1983) Esso System Visual Comparison

2 2

40 1 1

30

2

2

3

4

TAI

Spore Colour Index Fisher et al. (1980)

20 Thermal Alteration Index Staplin (1969)

2.0

10

50

100

Fig. 19. Correlation between SCI and TAI. Modified from Collins (1990). Data points of Smith (1983) are based on transmitted light spectra.

500 [Ma]

Duration of heating

Fig. 18. Karweil type diagram for calculation of maximum burial temperature. Modified from Cooper (1977) and Marshall (1990a).

μ-spectrophotometry (Marshall, 1990b, 1991; Milton, 1993). As the colour change in miospores is extremely subtle for a maturity below 1.0% VRr (Milton, 1993) this approach excellently complements the classic vitrinite reflectance and conodont colour alteration index measurements. At present, μ-spectrophotometry is the most promising method to quantitatively measure the miospore colour, provided that the complete spectrum is scanned and presented using the international colour system (Marshall, 1990b, 1991; Marshall and Yule, 1999; Milton, 1993). Quantitatively miospore colour measurements clearly show that below the oil birth line, miospore colour progressively changes from light yellow to yellow/orange and at about the main point of hydrocarbon generation a rapid series of colour changes occur from orange to brown. Beyond that stage, miospore colour (from dark brown to black) becomes rather insensitive to increasing temperature. Hence, the application of quantitative miospore colour measurements is especially relevant in the immature to mature stage and is an excellent method to identify the onset of hydrocarbon generation (Marshall and Yule, 1999). As the μ-spectrophotometry equipment is expensive and only available in few specialist laboratories, this approach has only been used in a limited number of studies (e.g. Hartkopf-Fröder et al., 2001, 2004; Milton, 1993; Oboh, 1992; Yule, 1998) since its introduction by Marshall (1990b, 1991). However, μ-spectrophotometry is the method of choice when quantitative and reproducible miospore colour data are needed. These data can be complemented by miospore exine reflectance (e.g. Bertrand et al., 1995; Littke, 1985; Marshall, 1991; Yule, 1998). The change of fluorescence colours has also been frequently applied as an indicator of thermal maturation (e.g. Senftle et al., 1993; Mao, 1994; Taylor et al., 1998). As maturity increases, the fluorescence of

the miospore exine shifts from green through yellow, orange and red to dark brown. Visible fluorescence of sporinite extinguishes at a vitrinite reflectance value of about 1.2–1.3%. Qualitative colour determination of fluorescence is, however, quite subjective. In order to circumvent this shortcoming, the spectral distribution of relative fluorescence intensities of miospores can be measured and various ratios can then be calculated from the acquired data. This well established approach was frequently used in exploration studies or to decipher regional trends in thermal maturity. 2.8.5.3. Chemical alteration of sporomorphs during diagenesis. The wall of sporomorphs, which is made up of sporopollenin, is extremely resistant and protects the protoplasm from unfavourable conditions until the moisture level is sufficient to allow fertilization by motile spermatozoids in the presence of water. The durability of the spore and pollen wall against e.g. physical abrasion, desiccation, UV-B radiation and biodegradation is the prerequisite for sporomorph dispersal and sexual reproduction in a subaerial environment. Sporopollenin represents one of the most resistant biopolymers and its insolubility in almost all solvents hampers the decipherment of the chemical structure. Two hundred years ago, Johann Friedrich John, a German chemist, created the term “Pollenin” after studying tulip pollen by ethanol extraction and subsequent precipitation experiments (John, 1814). He emphasized, as did Fritsche (1834), that “Pollenin” is insoluble in many chemicals. The numerous publications by Zetzsche and coworkers between 1928 and 1937 were a major progress in deciphering the chemical structure of the remarkably stable spore and pollen walls. They first used “Sporonin” to characterize the resistant compound of Lycopodium spores (Zetzsche and Huggler, 1928). When this team realized that the angiosperm “Pollenin” and the lycopsid “Sporonin” are very similar, they coined the term “Sporopollenin” (Zetzsche and Vicari, 1931b) which is still in use today. Zetzsche and Vicari (1931a, 1931b) determined a C/H ratio of 1 : 1.6 which is similar to that of terpenes and resulted in the chemical formula C90H127O12(OH)15. However, the insolubility of sporopollenin prevented a more detailed decoding of the structural formula. Shaw and Yeadon (1964) and Potonié and Rehnelt (1969) were probably the first to apply infrared spectroscopy to modern

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and fossil spores and pollen. This method is still an outstanding tool in the research on sporopollenin chemistry, although in present-day instruments a Fourier transform infrared spectrometer is coupled with an optical microscope which allows analysis of very small sample sizes at high spatial resolution (e.g. Steemans et al., 2010; Yule et al., 2000). A historical account of the beginning of research in sporopollenin chemistry is given by Shaw (1970). The more recent research on the chemical structure of sporopollenin has been summarized by de Leeuw et al. (2006). Modern analytical techniques such as 13C NMR spectroscopy (e.g. Bubert et al., 2002; Hemsley et al., 1996), FTIR and μ-FTIR spectroscopy (Bubert et al., 2002; Steemans et al., 2010; Watson et al., 2007; Yule, 1998; Yule et al., 2000), μ-Raman spectroscopy (Bernard et al., 2007, 2008; Ivleva et al., 2005; Schulte et al., 2008), Curie Point-, flash- or laser-pyrolysis–gas chromatography/mass spectrometry, partially with online derivatization (e.g. al Sandouk-Lincke et al., 2013; Blokker et al., 2006; van Bergen et al., 1993; Watson et al., 2007) and near edge X-ray absorption fine structure (NEXAFS) spectroscopy (Bernard et al., 2007, 2009) improved our understanding of the chemical structure of sporopollenin. Besides spectroscopical and pyrolytical methods, chemical degradation techniques such as acidic methanolysis have also been applied for structurechemical investigations (Bubert et al., 2002). Most pyrolytical studies have been carried out on megaspores while μ-FTIR spectroscopy can easily be done using a single miospore. In Holocene material two types of sporopollenin have been determined and it is possible that not all plant groups produce identical sporopollenin. One type is characterized by functionalized aromatic building blocks (pcoumaric and ferulic acid), the other consists predominantly of an aliphatic biopolymer of unknown structure. Fossil megaspores consist of both highly aliphatic (characterized by alkane/alkene doublets as main pyrolysis products) and aromatic moieties (de Leeuw et al., 2006), the latter indicated by e.g. acetophenone and 4-hydroxystyrene (= vinylphenol; a pyrolysis product of p-coumaric acid). Compared to modern sporopollenin, fossilized megaspores are enriched in aliphatic constituents. This might be due to i) selective preservation, i.e. aliphatic moieties being more resistant against diagenesis than aromatic compounds, or to ii) lipids which were present in the cell content or the surrounding sedimentary rock and which invaded or became attached to the spore wall sporopollenin by means of oxidative cross linking. Recent studies summarized by de Leeuw et al. (2006) suggest that the aliphatic moieties are in fact not indigenous to the original sporopollenin and that due to burial and diagenesis the primordial biopolymer is considerably altered even in thermally immature sedimentary rocks. On the contrary, Fraser et al. (2012) reported a high chemical similarity of thermally immature Pennsylvanian sporopollenin when compared to extant relatives. They concluded that land plant sporopollenin has remained stable from a chemical point of view since the Palaeozoic. Noteworthy, Watson et al. (2007) pointed to another parameter that seems to influence the chemical constitution of sporopollenin. Applying thermochemolysis GC/MS and μ-FTIR techniques they observed systematic changes of UV-absorbing moieties (p-coumaric and ferulic acid moieties) in recent spores from different altitudes. Consequently, they speculated that chemical modifications in fossil miospores might also reflect changing UV conditions e.g. by variations of the stratospheric ozone layer. Although many pyrolytic studies are devoted to the composition of modern and fossil sporopollenin, the chemical modifications due to the maturation process have been elucidated using this powerful method only to a very low extent. al Sandouk-Lincke et al. (2013) investigated Early Jurassic bisaccate pollen covering a thermal maturity series from immature to overmature. Pyrolysis and μ-FTIR analyses revealed two maturity-related trends, an increasing defunctionalization as well as an increasing aromatization. Hemsley et al. (1996, additional references therein) used 13C nuclear magnetic resonance spectroscopy to examine the alteration of functional groups but these studies are based on artificially heated acetolysed pine pollen. Yule (1998) and

Yule et al. (2000) applied μ-FTIR spectroscopy to analyse the chemical properties, i.e. the functional groups of sporopollenin during progressive thermal maturation. During the mature phase they observed the loss of a considerable portion of the aliphatic groups and an increasing formation of carbon-carbon double bonds associated with the formation of aromatic rings. With increasing maturity polycyclic aromatic units are formed (Yule et al., 2000). This observation is partially supported by results published by Watson et al. (2012). Artificial maturation also revealed increasing aromaticity of macromolecular sporopollenin. Additionally, a contemporary formation of polyalkyl hydrocarbons under anhydrous conditions has been attributed to maturationinitiated polymerization of hydrolysable components. Fraser et al. (2014) were able to demonstrate, also by artificial maturation, that a significant change of chemical properties started not until elevated temperatures have been reached linking substantial modification of sporopollenin to higher maturity. μ-FTIR analyses of artificially matured Lycopodium spores and Pennsylvanian megaspores showed strong similarities indicating that their heating experiments reflect the natural maturation process. 2.8.6. Chitinozoans 2.8.6.1. General remarks. The earliest chitinozoans, an extinct group of organic-walled microfossils and tentatively assigned to ontogenetic vesicles of unknown soft-bodied organisms (Paris and Nõlvak, 1999), have been recorded from the Cambrian Stage 5 (mid-Cambrian) of southern China (Shen et al., 2013). Recently, detailed reviews on chitinozoans have been published by Grahn and Paris (2011) and Servais et al. (2013). Poorly preserved and graphitized chitinozoans have even been extracted from high-grade metamorphic gneisses (Hanel et al., 1999). Schiffbauer et al. (2012) have shown that organic-walled microfossils do have a preservation potential in greenschist–amphibolite metasediments. 2.8.6.2. Colour alteration and reflectance of chitinozoans during diagenesis. Chitinozoans are frequently used as a maturity parameter, especially in marine pre-Devonian sedimentary rocks where vitrinite is rare or lacking (Supplementary material 2). With increasing maturity, under transmitted light chitinozoans display a colour change from amber through dark brown to black. Simultaneously, they become increasingly opaque and reflectant (Correia, 1967; Goodarzi et al., 1992). However, colour and translucency is also dependent on the wall thickness of the chitinozoan vesicles. As thickness varies considerably quantitative chitinozoan colour measurements have never been used. Instead, chitinozoan reflectance turned out to be a much more promising approach which is routinely applied in thermal maturity studies provided that vitrinite is absent or rare (e.g. Goodarzi, 1985; Obermajer et al., 1996; Tricker et al., 1992). Chitinozoan reflectance has been used in regional maturity and hydrocarbon exploration studies and in locating anomalies in organic matter reflectance due to mineralization (for references see Supplementary material 2). In polished blocks chitinozoans can be readily identified. However, usually they are rarely recorded in rock sections. To overcome this drawback kerogen concentrates are the first choice in order to enrich chitinozoans and other zoomorphs (for details see chapter 2.8.). With increasing thermal maturity chitinozoans show an increase in reflectance. They are optical isotropic and have thick, homogeneous vesicle walls which makes measurements quite easy and replicable. In samples rich in AOM where vitrinite reflectance is frequently suppressed, chitinozoan reflectance usually does do not show significant suppression (Tricker et al., 1992). However, a rare case of retardation in chitinozoan reflectance increase has been reported from the basemetal sulphide Polaris deposit (Arctic Canada). The factors causing the retardation are not yet deciphered in detail. Hydrogenation induced from algae-rich organic matter and oxygenation processes may result in suppressed chitinozoan reflectance (Héroux et al., 2000). Compared


to graptolites, scolecodonts and miospores, chitinozoan reflectance is always higher at a given rank (Bertrand, 1990a). Increase in reflectance of chitinozoans is linear over a large maturity range when calibrated to vitrinite reflectance. Correlations between chitinozoan reflectance and reflectance of other zoomorphs and between chitinozoan and vitrinite reflectance have been published in various studies (e.g. Bertrand, 1990a, 1993; Bertrand and Héroux, 1987; Bertrand and Malo, 2001, 2012; Tricker et al., 1992). 2.8.6.3. Chemical alteration of chitinozoans during diagenesis. Chitinozoans as well as scolecodonts can contribute to a lowering of HI values even if present in small amounts (Jacobson et al., 1988). The macromolecular composition of chitinozoans has been elucidated recently (Dutta et al., 2007; Jacob et al., 2007). The two studies are based on five samples of low thermal maturity (Tmax 420–433 °C, equivalent to VRc 0.45–0.6%). Dutta et al. (2007) used μ-FTIR spectroscopy and Curie Point-pyrolysis– gas chromatography/mass spectrometry, while Jacob et al. (2007) applied laser-pyrolysis–gas chromatography/mass spectrometry, μ-FTIR spectroscopy and laser μ-Raman spectroscopy. The macromolecules consist of both aliphatic and aromatic moieties, from which the latter moieties predominate over the aliphatic fraction. 1,2,3,4-Tetramethylbenzene, derived most probably from diaromatic moieties, is the most abundant pyrolysis product of the chitinozoans. Diagnostic pyrolysis products related to chitin have not been detected in both studies and hence, it is unlikely that the primordial biomacromolecules of chitinozoans prior to fossilization were made of chitin (Dutta et al., 2007; Jacob et al., 2007). An interesting study using laser μ-Raman spectroscopy has been presented by Roberts et al. (1995). The spectra of chitinozoans from 0.63 to 5.12% CR (= chitinozoan reflectance) show two broad bands of which the line-width of the 1600 cm−1 band (out of phase coupling of an in-plane C\\C stretching vibration) decreases with increasing rank. In the oil window this reduction is most pronounced proving that laser μ-Raman spectroscopy of chitinozoans can be used as a complementary method to characterize maturity in hydrocarbon exploration. Except for the laser μ-Raman spectroscopy study by Roberts et al. (1995) no data are available on the alteration of the chitinozoan organic structure with increasing rank. 2.8.7. Scolecodonts 2.8.7.1. General remarks. Scolecodonts represent fossilized jaw elements of polychaetous annelids of the orders Eunicida and Phyllodocida. Complete jaw apparatuses are rarely preserved in the fossil record. They are complex and composed of numerous, up to more than 20 jaw elements of different size and shape. Most Palaeozoic scolecodonts belong to the Eunicida which have also some descendants in the Mesozoic and Cenozoic (Paxton, 2009; Szaniawski, 1996). Of the Phyllodocida which appear in the Dzhulfian (Late Permian) (Nakrem et al., 2001) two families are present in the fossil record, the Glyceridae and Goniadidae. The earliest scolecodonts have been reported from the latest Cambrian (latest Furongian) of Newfoundland (H.S. Williams in Hints and Eriksson, 2007), but they were most diverse and numerous during the Mid/Late Ordovician, Silurian and Devonian. In the Carboniferous abundance and diversity decreased and in the Permian scolecodonts suffered a major extinction. Postpalaeozoic scolecodonts resemble modern forms, are relatively scarce, and taxonomic diversity is much lower compared to early Palaeozoic assemblages (Courtinat, 1990; Szaniawski, 1996). Scolecodonts are usually of quite limited value in biostratigraphy. As their main radiation occurred in the Darriwilian (Mid Ordovician) and assemblages were diverse until the Devonian, scolecodonts have some biostratigraphic potential at least on a regional scale provided that large collections are available (e.g. Bergman, 1989; Eriksson and Bergman, 2003; Eriksson et al., 2004). A drawback for their use in establishing even a relatively coarsely resolved biozonation is the facies dependence of scolecodonts. Usually, they are more abundant in shallow shelf settings predominantly in argillaceous limestones. In reef or deep water facies

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they are much rarer. However, some species were specialized to reefal and lagoonal environments. Some polychaete worms can survive under dysoxic conditions but in oxygen deficient and deep water environments the size of scolecodont specimens is relatively small (e.g. Eriksson et al., 2004; Tyson, 1995). Scolecodonts are within the size range of 0.1 to 4 mm (Szaniawski, 1996) and the concentration usually does not exceed a few specimens/gramme (Eriksson et al., 2004; Hints and Eriksson, 2007). However, up to 1500–2000 specimens/gramme have been reported on some rare occasions (Head, 1993). As scolecodonts are resistant to most acids they are found in palynological preparations and in the residue after processing for phosphatic fossils, e.g. conodonts. The reflectance of scolecodonts in polished blocks has been used in thermal maturity studies (see below). However, as they occur in low concentrations in sedimentary rocks it can be difficult to achieve a sufficient number of readings. In this case using kerogen concentrates (e.g. Bertrand and Malo, 2012; Bertrand et al., 1985) or applying standard palaeontological techniques (i.e. dissolving sedimentary rock, handpicking, and embedding the scolecodont concentrate in epoxy resin for polishing, e.g. Goodarzi and Higgins, 1987) will improve the data base with reasonable efforts. Although most of the jaw elements are very characteristic and therefore difficult to misinterpret, care must be taken not to confuse scolecodonts (especially the maxillae I of some Eunicida and the jaws of Phyllodocida) with cephalopod arm hooks (Kulicki and Szaniawski, 1972). 2.8.7.2. Colour alteration and reflectance of scolecodonts during diagenesis. It is well known that scolecodont colour under transmitted light changes from yellow and reddish brown in sedimentary rocks of low thermal maturity through black and grey with increasing palaeotemperature (Dorning, 1986; Goodarzi and Higgins, 1987; Goodarzi et al., 1992). Scolecodonts have even been reported from slightly metamorphosed sedimentary rocks (CAI 5–5.5 equivalent to burial temperature of over 300 °C) but they are badly preserved and the organic matter commonly replaced by microgranular pyrite (Suttner and Hints, 2010) which prevents colour and reflectance analyses. As scolecodonts vary in thickness they are translucent in the thinnest and opaque in the thickest parts and hence, quantitative colour measurements are difficult to obtain although the general colour impression reveals a rough indication of the thermal maturation. In addition, colour seems to be somewhat taxon-dependant (Bergman, 1989, p. 31). Instead of colour variations, scolecodont reflectance has been used as a quantitative maturity parameter (Supplementary material 2). Because of the inner structural differentiation of scolecodonts in layers with pore-canals of various diameter and orientation (e.g. Szaniawski, 1996) the measurements must always be taken from homogenous parts. These ultrastructural details can be observed in polished scolecodont specimens and areas with pores-canals or cavities must be avoided. Optimal surfaces are the scolecodont tips and the outermost compact layer (Bertrand, 1987, p. 226). Reflectance of scolecodonts is well studied up to a vitrinite reflectance of approx. 2.5% VRr (CAI approx. 3.5–4.0). They are isotropic at least up to this maturity level (Goodarzi et al., 1992). In general, scolecodont reflectance increases with increasing maturity. Although scolecodonts are less reflecting than vitrinite, chitinozoans and graptolites, the optical properties of all groups follow a similar trend (Bertrand, 1990a; Bertrand and Héroux, 1987; Goodarzi and Higgins, 1987; Obermajer et al., 1999b). To convert scolecodont reflectance into reflectance of other zoomorphs or into vitrinite reflectance, equations published by e.g. Bertrand (1990a, 1993). Bertrand and Héroux (1987) and Bertrand and Malo (2001, 2012) can be consulted. 2.8.7.3. Chemical composition of scolecodonts. While the soft parts of polychaetes are rarely preserved, scolecodonts have a high fossilization potential. Not much is known about the biogeochemical composition of modern and especially of fossil polychaete jaws. Traditionally,

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scolecodonts have been regarded as being composed of chitin and this statement has been adopted in current textbooks (e.g. Traverse, 2008). However, since the early 1960s several studies proved that in modern jaws amino acids such as glycine and histidine are dominant and that metals (e.g. copper, iron, zinc) are particularly highly concentrated in the tips (e.g. Eriksson and Elfman, 2000; Jeuniaux, 1963; Lichtenegger et al., 2002; Voss-Foucart et al., 1973; for more references on modern polychaete jaw chemistry see Dutta et al., 2010). Scolecodonts, the fossilized jaw elements, can presumably be classified as Type III kerogen (Tyson, 1995) and also have metal enrichments in areas which had been more exposed to abrasion in the living animal (Eriksson and Elfman, 2000). According to μ-FTIR spectroscopy, Curie Point-pyrolysis–gas chromatography/mass spectrometry and TMAH assisted thermochemolysis–gas chromatography/mass spectrometry analyses, scolecodonts from samples of Upper Ordovician–Middle Devonian immature to very early mature sedimentary rocks (Tmax 416–434 °C, corresponding to VRc 0.45–0.6%) are composed of both aliphatic and aromatic moieties. The major pyrolysis products from these scolecodonts are aromatic compounds (e.g. alkylbenzenes, alkylnaphthalenes and alkylphenols) and aliphatic hydrocarbons (e.g. a homologous series of n-alkenes and n-alkanes). No protein/amino acid derived compounds have been recognized which is attributed to diagenetic alteration (Dutta et al., 2010). Gupta and Briggs (2011, p. 205) found a dominant aromatic composition in scolecodonts from the Upper Ordovician of Cincinnati but no thermal maturity data are provided. In a comparative study of Curie Point-pyrolysis–gas chromatography/ mass spectrometry and laser-pyrolysis–gas chromatography/mass spectrometry, Late Devonian–Mississippian scolecodonts of high thermal maturity (approx. 3.4% Rc) revealed a predominance of aromatics (benzene, toluene, xylene, naphthalene, methylated naphthalenes, styrene), low amounts of aliphatic hydrocarbons in the range of n-C6 to n-C19 and few oxygen containing compounds (phenol, cresol and benzaldehyde). Curie Point temperatures lower than 764 °C provided low pyrolysis yields implying that at this high level of thermal maturity the scolecodonts consist of an organic structure resistant to thermal break-down (al Sandouk-Lincke et al., 2014). μ-FTIR data of immature to very early mature scolecodonts show that they are more aromatic than aliphatic and hence the I1 ratio which compares the relative abundance of aliphatic and aromatic components is low and similar to chitinozoans. Additionally, the CH2/CH3 ratio is also low, expressing short and more branched aliphatic chains. A- and C-factors which represent relative intensities of the aliphatic and C_O groups, respectively, are both low elucidating that aromatic compounds prevail over aliphatic components. In an I1 vs. CH2/CH3 and a C-factor vs. A-factor diagram scolecodonts plot near chitinozoans (Dutta et al., 2013). Unlike the above-mentioned studies, μ-FTIR spectra of scolecodonts analysed by Nowaczewski (2011) and Olcott Marshall et al. (2013) reveal an aliphatic polyether macromolecule. Samples for this study were collected from near the Devonian–Carboniferous boundary in the Arbuckle Mountains. Conodonts recovered from the same samples are brown in colour (CAI 2) which is equivalent to 0.7–1.5% VRr (mature to postmature stage with regard to hydrocarbon generation). At present, it is difficult to identify the reasons for the different results. Olcott Marshall et al. (2013) suggest that the aliphatic scolecodonts may be better preserved chemically and may reflect the primary composition of the scolecodonts more closely or that scolecodonts may have different initial biopolymer composition. Other possible causes for the contradictory results could be differences in thermal maturity, host sediment or sample processing. The application of acetic acid, bleach and mineral spirit solution as used in the sample preparation by Nowaczewski (2011) and Olcott Marshall et al. (2013) may be problematic as it is unknown whether these reagents have effects on the chemical composition of palynomorphs. Oxidizing agents in particular should always be avoided while the commonly applied hydrochloric/hydrofluoric acid digestion as employed by Dutta et al. (2010, 2013) is thought not to significantly influence the chemical composition of

fossil organic material as long as it is not very immature (for further discussion see van Bergen, 1999). Systematic studies on scolecodont macromolecular alteration due to increasing thermal maturity have not yet been undertaken even though scolecodonts do have an influence on hydrocarbon source rock quality (Jacobson et al., 1988). 2.8.8. Melanosclerites and hydroids Melanosclerites are occasionally encountered in palynological residues. The earliest melanosclerites have been reported from lower Cambrian rocks of the East European Platform. They have not yet been found in post-Frasnian strata but as this enigmatic palynomorph group seems to have only limited stratigraphical and palaeoecological potential, they may have been overlooked or neglected (Cashman, 1996; Winchester-Seeto and McIlroy, 2006). Melanosclerites seem to be restricted to marine environments. Although they are highly variable in shape, rod- to bell-shaped melanosclerites with bulbous, tapered or blunt ends prevail. They range from approx. 100 μm to several mm in size. Their biological affinity is still a matter of debate. Melanosclerites have been assigned to algae but in most studies they have been interpreted as larval or early polyp stages of various Medusozoa (Cnidaria) (Cashman, 1992, 1996; Winchester-Seeto and McIlroy, 2006). In fact, many melanosclerites show strong resemblance with hydroids and in some samples they occur together with hydrozoan periderms (Cashman, 1996). While melanosclerites are mainly recognized in palynological preparations, hydroid remains are consistently observed in polished blocks of marine sedimentary rocks of Ordovician to Devonian age (Bertrand and Malo, 2012; Héroux et al., 2000). The earliest record of a hydroid cnidarian colony comes from the Early Ordovician of China (Baliński et al., 2014). In polished blocks hydroids are generally less often recorded than graptolites, chitinozoans or scolecodonts (e.g. Bertrand, 1987; Héroux et al., 2000) but in proximal marine sedimentary settings they can be a major constituent of the zooclast/zoomorph fraction (Bertrand, 1990b, 1991; Bertrand and Malo, 2012). In transmitted light hydroids are dark brown and in reflected light they are grey and isotropic (Bertrand, 1987). The reflectance values are probably not influenced by the lithology of the host rock (Bertrand, 1993). Reflectance of hydroids has been used as a supplementary parameter in some maturation studies but reflectance data of graptolites, chitinozoans and scolecodonts are more often employed as a substitute for vitrinite (e.g. Bertrand, 1987, 1990b; Bertrand and Malo, 2005, 2012; Bertrand et al., 2003; Héroux et al., 1996, 2000). This may be because hydroids are generally less abundant than other zooclasts and zoomorphs. In addition, their walls are thinner and more delicate than those of graptolites (for photographs of hydroids in polished mounts of kerogen concentrates see e.g. Bertrand, 1987, Plate 7.9 and 8.6; Bertrand and Malo, 2012, Plate 1–2). Fragments of hydroids may therefore have been summarized as undifferentiated zooclasts. Studies in which hydroid reflectance was used along with other zooclasts and zoomorphs are listed in the Supplementary material 2. Following simple solubility experiments Eisenack (1963) concluded that melanosclerites consist of chitinous matter mixed with various proteins. Cashman (1992) suggested a “pseudochitinous” composition similar to e.g. chitinozoans, scolecodonts or graptolites while Cashman (1996) pointed out that the organic material was previously summarily identified as chitinous but that a validation based on modern techniques is lacking. So far nothing has changed regarding this statement. Fossil hydroids are believed to consist of chitin (Bertrand, 1987) but modern organic geochemical analyses are lacking. 2.8.9. Arthropod cuticles 2.8.9.1. General remarks. Most arthropods are not microfossils and consequently small fragments should not be referred to as palynomorphs but rather as zooclasts. However, particularly in Palaeozoic samples


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small fragments of arthropod cuticles are regularly found in palynological preparations. The cuticle is the multilayered outer part of the arthropod exoskeleton. Even complete compound eyes, chelicerae and appendages (e.g. pedipalps) can be preserved and larger fragments can show characteristic features such as pores, sensory hairs and spines. Identification of the fragments on genus or family level may be possible but smaller early Palaeozoic remains can be difficult to distinguish from plant cuticles particularly when dealing with highly carbonized fragments in which characteristic ornaments are destroyed (Gensel et al., 1990). In Palaeozoic sedimentary rocks most of the dispersed arthropod cuticles can be assigned to eurypterids or scorpions (Braun, 1997; Haug et al., 2014; Miller, 1996) while from the Mesozoic and Cenozoic a variety of other arthropods has been reported including the organic linings of ostracod carapaces and insects (Bartram et al., 1987; Tyson, 1995, 212). Arthropod cuticles have also been isolated by maceration from coals (Bartram et al., 1987) and while these remains can make up 20–30% of all cuticle fragments (e.g. Wilson and Hoffmeister, 1956) they have not yet been closely examined in coal petrological studies. 2.8.9.2. Colour alteration of arthropod cuticles during diagenesis. In transmitted-light microscopy the cuticles are translucent and depending on thermal maturity light yellow (e.g. in the immature Miocene Rhenish brown coal) to grey and black in the anthracite stage (Fig. 20). Goodarzi (1984b) described optical properties of arthropod cuticles recovered from sub-bituminous coal. In reflected light these cuticles are dark grey and isotropic, low reflecting and have a yellow to dark orange fluorescence. Low reflectance and strong fluorescence make arthropod cuticles of this coalification stage optically similar to liptinites, especially to resinites (Goodarzi, 1984b). Strong fluorescence has also been reported from some fossil decapod carapaces (e.g. Haug et al., 2009; Huang et al., 2013). As has been pointed out by e.g. Bartram et al. (1987), Braun (1997) and Dorning (1986) optical properties of arthropod cuticles may still be an unexplored tool for roughly determining thermal maturity levels. In coals and in thin mudstone layers intercalated within coal seams, arthropod cuticles seem to be particularly associated with charcoalified plant remains and fusinite-rich coals (Bartram et al., 1987). 2.8.9.3. Chemical composition of arthropod cuticles during diagenesis. Modifications in the chemical composition of arthropod cuticles over a wide range of thermal maturity have not yet been studied systematically. Modern arthropod cuticles are multilayered, covered by a waxy layer and constructed of a composite material incorporating chitin fibres embedded within a protein matrix. Parts of the cuticle may be sclerotized and in numerous aquatic arthropods the cuticle is biomineralized with calcium carbonate (Gupta, 2011; Gupta and Summons, 2011). As chitin is readily degraded by oxidation, hydrolysis and chitinolytic microorganisms it is rarely found in the geological record and even fossil arthropod cuticles of Cenozoic age and of very low thermal alteration are mainly composed of an aliphatic geopolymer while traces of chitin could only be detected by employing pyrolysis–gas chromatography/ mass spectrometry (see Gupta and Summons (2011) for details). Applying this method to eurypterid cuticles from early mature sedimentary rocks of Silurian (Ludlow, Pridoli) and Pennsylvanian age revealed that fatty acyl moieties are the most abundant compounds while chitin is absent. At slightly higher thermal maturity the presence of phenols and polyaromatic compounds indicates a higher degree of aromatization. Correspondingly, the Raman spectra show an increase in ordered carbonaceous matter (Gupta et al., 2007; Stankiewicz et al., 1998). However, pyrolysis–gas chromatography/mass spectrometry may not always be the first choice to detect remnants of the degraded chitin– protein complex. Instead, based on X-ray absorption near edge structure spectromicroscopy analyses (XANES) and some additional techniques Ehrlich et al. (2013) discovered chitin in a demosponge from

Fig. 20. Arthropod remains in transmitted light microscopy showing darkening with increasing thermal maturity. (a) Sewer construction site Mühlenbach, Essen-Barbeck, Germany, Quaternary, thermally unaltered; (b) slagheap Rudolfschacht, Ibbenbüren, Germany, Duckmantian (c) Cemex quarry, Piesberg, Germany, Asturian.

the middle Cambrian Burgess Shale. Applying XANES spectroscopy Cody et al. (2011) were able to trace high amounts of the vestigial chitin–protein complex in a Pennsylvanian scorpion cuticle (0.39% VRmax in oil) and a Pridoli eurypterid cuticle, the latter recovered from the

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same locality which already provided the material analysed by Gupta et al. (2007). Besides the chitin–protein relict, XANES spectroscopy revealed an increase in aliphatic carbon which is in accordance with results from the pyrolysis–gas chromatography/mass spectrometry studies. Hence, the primordial chitin–protein complex is already considerably altered at an early stage of diagenesis and partly replaced by a predominantly aliphatic composition. This is attributed to in situ transformation and incorporation of labile alkyl compounds to form the highly aliphatic geopolymer (Gupta and Summons, 2011). In fact, confined heating of arthropod cuticles confirmed cuticular lipids as the source for the aliphatic signature in fossil cuticles (Gupta, 2011). 3. Discussion Vitrinite reflectance is the most widely applied optical maturity parameter which can be applied to all post-Silurian sedimentary rocks containing huminite/vitrinite particles. Identification of these particles is difficult, where different vitrinite populations (autochthonous and resedimented) exist or where particles are very small or rare. Also, differences in vitrinite reflectance for different lithologies have been noted (Bostick and Foster, 1975; Jasper et al., 2009; Scheidt and Littke, 1989; see discussion in Suárez-Ruiz et al., 2012) which are probably due to different chemistry of vitrinite precursor material or early diagenetic transformation. In addition, differences in thermal conductivities may play a role which can be more than 10 times higher in sandstones (and salt) as compared to coal (about 0.4 W/(m*K)). These differences can only be neglected, if rocks from narrow depth intervals are compared, i.e. some metres or tens of metres apart. For rocks that have reached a high maturation stage corresponding to VRr N 2.0%, the emerging anisotropy has to be taken into account. This anisotropy depends not only on temperature but also on stress and strain which differs in different types of rocks. In a recent study, Bruns and Littke (2015) could demonstrate that rotational reflectance of vitrinite provides a powerful tool to study rock strain. On the other hand, for rocks close to metamorphism, calculation of maximum burial temperatures from vitrinite reflectance becomes more uncertain than for rocks in the diagenetic realm.

Solid bitumen reflectance is another widely used tool, because many rocks contain only little or no vitrinite but abundant solid bitumen. Examples are petroleum source rocks, but also reservoir rocks. In carbonate reservoir rocks vitrinites tend to be extremely rare or absent, but solid bitumen can occur in large quantities. Also, fractures and faults can be partly filled with solid bitumen as residue of a former oil filling. One problem is related to the fact that in some of these reservoir rocks and in particular in fractures, different populations of solid bitumen occur, which show high and low reflectance probably representing different pulses of petroleum charge (Schoenherr et al., 2007). Also solid bitumen in fractures can be altered by hot fluids, which do not affect the surrounding rock matrix. Vitrinite reflectance and conodont colour alteration are among the most widely used maturity parameters. For low-grade diagenesis studies (up to about 1.5% VRr) vitrinite reflectance is more useful in comparison to conodonts as vitrinite is more sensitive. On the other hand conodonts are very useful at a high diagenetic stage up to metamorphism. Thus, both maturity parameters complement one another. Using the CAI values has many advantages, but also some pitfalls (see chapter 2.2). The application of CAI is a rapid and inexpensive method requiring only standard laboratory techniques and is very useful from anchimetamorphism to high grade metamorphism (CAI 5–CAI 8) and contact metamorphism. Thus, conodonts have been used to determine thermal aureoles related to igneous rocks (e.g. Armstrong and Strens, 1987; Königshof, 1991; Nicoll, 1981; Swift, 1993; Wiederer et al., 2002). Some authors tried to quantify conodont colour alteration data using infrared spectroscopy (Nöth and Richter, 1992), electron spin resonance spectroscopy (Belka et al., 1987), fluorescence spectrometry (Mastalerz et al., 1992), pyrolysis–gas chromatography/mass spectrometry (Bustin et al., 1992; Kemp, 2002; Marshall et al., 1999), X-ray photoelectron spectroscopy and Fourier transform infrared emission spectroscopy (e.g. Kemp, 2002; Marshall et al., 1999), spectral reflectance (Deaton et al., 1996), colour and digital image analysis (Bábek et al., 2008; Helsen et al., 1995) and recently Voldman et al. (2010) provided a computer model (EasyCAI) in order to estimate a time-temperature history for a given stratigraphic level. Geochemical studies shed more light in the complex organic geochemistry of conodont elements. The method

Fig. 21. Charred (grey) and uncharred (yellow) miospores. (a)–(d) Trilete miospores from middle Cretaceous karst infillings, Rhenish Massif, Germany; (e)–(f) bisaccate pollen from the Bolzano Volcanic Complex, Permian, Southern Alps.


a)

base of Gröden Sandstone CIELAB L*

60

CIELAB b*

100

80

40 60 20 40

CIELAB a* 0 20

40

CIELAB Chroma 20

-10

0

b)

20

40

Bolzano Volcanic Complex CIELAB L*

60

CIELAB b*

100

80

40 60 20 40

CIELAB a* 0 20

40

CIELAB Chroma 20

-10

0

c)

20

40

Bolzano Volcanic Complex CIELAB L*

60

CIELAB b*

100

80

40 60 20 40

107

used by Deaton et al. (1996) was successful and provided exact data, but it requires a large number of conodont elements per sample. Helsen et al. (1995) measured intensities of red, green, and blue colour components of reflected light (RGB mode) from selected small areas of polished conodont elements. Data show a good correlation between RGB intensities to CAI, but the method is also partly subjective due to the selection of measured areas of the conodont element. Furthermore, this method has a potential error of approximately half a CAI index for a given determination which is the same potential error arising from visual comparison with the conodont colour standard (Brime et al., 2003). Few studies have been published dealing with correlation with other optical maturity parameters, such as scolecodonts (Goodarzi et al., 1992). Scolecodonts are isotropic up to at least a CAI of 3.5 and their reflectance shows a good correlation with CAI. Among optical maturity parameters CAI values and vitrinite and huminite reflectance are most commonly used in marine carbonates and siliciclastics, respectively. The overall problem is the less sensitive evaluation of CAI temperatures in comparison to vitrinite reflectance. Even if some comparative studies of CAI and IC or Kübler Index (KI) and CAI and vitrinite reflectance show the same metamorphic regime (e.g. Brime et al., 2008; Buggisch, 1986; García-López et al., 1997; Helsen and Königshof, 1994; Kisch, 1990; Königshof, 1992; Kovács and Árkai, 1987; Wiederer et al., 2002), some questions remain. It is uncertain, whether the transition between diagenesis and anchizone (180–230 °C, e.g. Frey and Robinson, 1999; Kisch, 1990) corresponds to CAI 4 or CAI 5. García-López et al. (1997) suggested correlating the transition with CAI 4. Others, such as Kovács and Árkai (1987), correlate the transition from diagenesis to anchizone with CAI 5. The transition between anchizone and epizone occurs at about CAI 5.5 (García-López et al., 1997). In a further regional study on CAI and KI by Brime et al. (2003) the authors suggest that the range of CAI values inferred for the anchizone should be enlarged to include CAI values up to 6. Furthermore, it seems obvious that kinetics of illite evolution was retarded in respect to organic matter changes of conodont elements (Árkai et al., 2002; Brime et al., 2001; Hillier et al., 1995). Rejebian et al. (1987) have published an Arrhenius diagram, based on heating experiments, which models the evolution of CAI for different maximum temperatures and duration of heating. Their temperature estimations are obviously too high as assumed by some authors (e.g. García-López et al., 1997; Wiederer et al., 2002). Similar results have been published by Sudar and Kovács (2006) on regionally metamorphosed rocks from Serbia and Hungary. According to these authors mineral paragenesis, illite crystallinity as well as vitrinite reflectance data above the diagenetic zone indicate considerably lower temperatures than could be deduced from CAI data (Rejebian et al., 1987). Several maturation studies have been used to determine thermal aureols related to igneous rocks (e.g. Brime et al., 2003; Burnett et al., 1988; Königshof, 1991; Nicoll, 1981; Rejebian et al., 1987; Swift, 1993; Wiederer et al., 2002) but little is known why conodonts exhibit a wide range of CAI values within a single sample which is a characteristic feature of contact metamorphosed conodonts. However, the range of up to four CAI values within a single sample does not reflect temperature differences. Some assumptions have been made, e.g. variations of conodont colour may be related to migration patterns within the conodonts or along microfractures (Rejebian et al., 1987), or may be also subject of staining effects as suggested by Legall et al. (1981). Anyhow, these statements are more or less assumptions, a cross-validation with geochemical proxies is missing so far.

CIELAB a* 0 20 -10

40

CIELAB Chroma 20 0

20

40

Fig. 22. CIELAB chromaticity and colour tone diagrams of bisaccate pollen from the Permian of the Southern Alps. (a) From mudstone at the base of the Gröden Sandstone; (b) from chert layer in the Bolzano Volcanic Complex. These miospores display weaker colours (lower b* and lower Chroma values) than those from the Gröden Sandstone (see also Fig. 13b); (c) from chert layer in the Bolzano Volcanic Complex. The plot shows two well separated assemblages. Miospores with darker colours (lower b* and lower L* values) are charred. Modified from Hartkopf-Fröder et al. (2001).

108


100

Relative Intensity

Staßfurt Fm

Röt Fm Staßfurt Fm Leine Fm

50

0 430

500

600

700 [ nm]

Wave Length salt

mudstone

Fig. 23. Comparison of spectral distribution of fluorescence intensities of miospores derived from Zechstein/Buntsandstein salt and mudstone, western Germany. Modified from Ecke (1986).

Deformation of conodonts is well known in low grade metamorphosed rocks (e.g. Königshof, 1992; Kovács and Árkai, 1987) although not as a regular phenomenon. Based on studies in the Montagne Noire (Wiederer et al., 2002) it became evident that the temperature level attained at CAI 5 defines the brittle–ductile boundary of fine-grained apatite for unknown lithostatic pressures and strain rates prevailing. Deformation structures do not seem to influence the colour of conodonts in general because deformation occurs also in conodonts with low CAI values. The influence of diagenetic effects as well as hydrothermal activity on the conodont colour has been confirmed in many studies (e.g. Belka, 1993a; Legall et al., 1981; Rejebian et al., 1987). For instance, the relationship between redistribution of silica in carbonates and recrystallization of conodont apatite is obvious, but factors governing these patterns in detail are not yet well-known. Although the colour of sporomorphs is one of the best established maturity parameters, there are some pitfalls one should be careful to avoid. In an assemblage of yellow/orange and brown miospores the darker coloured specimens are usually interpreted to be reworked from older strata. However, during the maturity path a rapid transition from yellow/orange to brown colours occurs around the main point of hydrocarbon generation. These assemblages therefore do not always consist of a mixture of in situ and reworked miospores but are indicative for a specific maturity interval. In general, miospores offer a

Fig. 24. Retarded miospore colour change due to different palynofacies; lowermost Permian, Saar-Nahe Basin, Germany. Tmax values are nearly identical and indicate an immature to very early mature stage while miospore colour is clearly different. (a)–(b) Kerogen rich in AOM with bright yellow coloured monosaccate miospores, borehole Münsterappel I, 114.46 m, Tmax 433 °C; (c)–(d) kerogen rich in woody material with orange coloured monosaccate miospores, borehole Münsterappel I, 111.12 m, Tmax 431 °C.


2.58 23.03

ichthyoliths

conchostracans

ostracods

foraminifera

graptolites

conodonts

amorphous organic matter

arthropod cuticles

melanosclerites/hydroids

Sandstone which overlies the Bolzano Volcanic Complex (Fig. 22a, b). In addition, fluorescence properties of the Gröden Sandstone miospores point to higher thermal maturity compared to the assemblage from the older Bolzano Volcanic Complex (Fig. 13b). Pronounced differences in fluorescence spectra of miospores from Zechstein mudstones and salt were also observed by Ecke (1986). The wavelength of the spectral peak of miospores from the Zechstein salt is around 550 nm while the wavelength of the spectral maximum of miospores from mudstones is near 700 nm (Fig. 23). Dependent on the kerogen type, miospore colour change may be retarded (Fig. 24). Miospores from AOM-rich kerogen show lighter colours (about one to two points less on the SCS scale) than miospores from kerogen which is mainly composed of woody material (e.g. Hillier and Marshall, 1992; Marshall et al., 1985). A particularly interesting example of miospore colour deviating from the normal maturity

scolecodonts

chitinozoa

prasinophycean algae

dinoflagellates

acritarchs

sporomorphs

vascular plants

System/Period

Erathem/Era

Cenozoic

numerical age [Ma]

biostratigraphic control and differentiation between genuine and reworked specimens should be possible. Charring of miospores by wildfires can also result in assemblages comprising very dark coloured and much lighter coloured miospores (Hartkopf-Fröder et al., 2001). Frequently, the charred miospores are grey to black, wrinkled and badly preserved (Fig. 21). In chromaticity and colour tone plots, charred miospores from the Bolzano Volcanic Complex (Permian, Southern Alps) are easily to distinguish from the non-charred specimens (Fig. 22c). Various studies demonstrated that miospore colour is not only related to thermal maturity but can be strongly influenced by the petrography of the host sediment and the kerogen composition. Hartkopf-Fröder et al. (2001) reported a distinct colour difference between lighter coloured miospores from chert layers of the Bolzano Volcanic Complex and darker coloured miospores from mudstones of the basal Gröden

109

Quaternary Neogene

Paleogene

66.0

145.0

Mesozoic

Cretaceous

8 Jurassic

201.3 Triassic 252.17 Permian 298.9

358.9

419.2 443.8

Palaeozoic

Carboniferous

10 6

Devonian

2

Silurian Ordovician

11

5 4 3

1 9

485.4 Cambrian 541.0

7

Fig. 25. Stratigraphic range and important evolutionary steps of fossil groups which are useful in thermal maturity studies. 1 = earliest bryophyte-like plants, MidOrdovician; 2 = earliest tree-sized plants, near Eifelian–Givetian boundary and earliest seed plants, Mid Givetian; 3 = earliest trilete miospores, Katian–Hirnantian; 4 = earliest common trilete miospores, Llandovery; 5 = earliest megaspores, Emsian; 6 = earliest saccate (protosaccate) miospores, late Frasnian; 7 = earliest putative dinoflagellate cysts, Mesoproterozoic; 8 = earliest terrestrial dinoflagellates, Portlandian (Tithonian); 9 = earliest planktic graptoloids, basal Tremadocian; 10 = earliest terrestrial ostracods, Mississippian; 11 = earliest non-marine ichthyoliths, late Famennian. Note that the modern pterobranch Rhabdopleura is regarded as an extant graptolite. Time scale and numerical ages follow the International Chronostratigraphic Chart, version 2015/01; http://www.stratigraphy.org/ICSchart/ChronostratChart2015-01.pdf.

110


thermal maturity analyses (Fig. 25). Aside from organic matter, fission tracks (in apatite and zircon), fluid inclusion homogenization temperatures and clay mineral based parameters such as illite crystallinity provide additional and important information on temperature evolution. Maturation of sedimentary organic matter is of greatest importance to understand petroleum systems, i.e. the dynamics of oil and natural gas generation in the deep subsurface. In combination with numerical basin modelling tools, these parameters allow to reconstruct the thermal history of sedimentary rocks providing important constraints for tectonic evolution. Vitrinite, solid bitumen and graptolite reflectance are the most commonly used parameters measured in reflected light. At low levels of maturation the original precursor material (of vitrinite) has some influence on the reflectance. This fact can be easily deduced from reflectance measurements on coals, in which vitrinite is usually the predominant material. Even at hard coal (bituminous coal) rank, reflectance of

path has been reported by van de Schootbrugge et al. (2009). Near the Triassic–Jurassic boundary miospores are much darker than from strata below or above. This miospore ‘dark zone’ is well known from the Danish–German Basin and possibly also from offshore eastern Canada (Weston et al., 2012). A natural darkening caused by sulphuric acid deposition in connection with flood basalt volcanism is assumed (van de Schootbrugge et al., 2009). 4. Conclusion and outlook Organic particles undergo systematic changes of optical and chemical properties during burial within sedimentary basins which are irreversible. Thus, numerous organic matter maturity parameters have been established reflecting palaeotemperature evolution and in particular maximum palaeotemperature. Depending on the stratigraphic position of the sampled interval various fossil groups can be used in

Coal and hydrocarbon stages and correlation of important optical thermal maturity parameters 1 Vitrinite VR r

Coal Rank

Hydrocarbons Stages and Tmax

[%]

Types

Zones of Hydrocarbon Generation and Destruction

CAI

2

3

4

TCI

AAI

5

AAI

6 SCI

1981

lignite

0.8

immature

biogenic dry gas

590

subbitum. B A

high B volatile bitum.

1990

2012

GR r [%] 2013

14

CR r

CR r

CR r

[%]

[%]

[%]

1990

1992

2012

1990

15 SR r [%]

16 SR r [%]

17 HR r [%] 2012

2012

1

1

2

2

3

3

2.2

600

0.5 0.5

2.4 5

4

1

2.0

6

5

mature

oil

oil 610

2.8

0.5

2 3

2.5

8

471 °C

wet gas

postmature

condensate

0.5

1.0 3.0

1.0 6

1.0

1.0

1.5

630 4

dry gas

3

1.5

650

9 5

3.5

4.0

1.5

2,0

overmature

thermogenic dry gas

2,0

2.0

2,0

2.5

2.5 4.0

10

7

3.0

2.5

3.0

2.5

2.5 3.0

3.0

1.5

2,0 2.5

670

10

1.5 2,0

9

3.0

1.5

1.0

2,0

660

2.5

1.0

1.5 1.5

3.0

640

1.0

1.0

1.5

4

metaanthrac.

1.0

7

620

semianthrac.

5.0

0.5

0.5

2.6

(515 °C)

anthracite

0.5

6

8

2.0

0.5

7

2 medium volatile bitum.

0.5

4

424 °C

A

low vol. bit.

4.0

[%]

13

4

1.5

3.0

GR r

[%]

12

3

5

1.0

1.35

GR r

11

1.5

C

0.9

1.2

1.0

1

C

0.6 0.7

SCS AMC/ ICM

10

1 2

0.5

9

1998

peat

0.4

TAI

8

[nm]

0.2

0.3

7

3.0

2,0 2.5 3.0

2,0

2.5

2.5 3.0

3.0

5

Fig. 26. Coal stages, phases of oil and gas generation and correlation between optical thermal maturity parameters. Column 1: Correlation between Colour Alteration Index (CAI) and vitrinite reflectance (VRr) from Taylor et al. (1998, Fig. 3.40). Column 2: Correlation between Transmittance Color Index (TCI) and vitrinite reflectance (VRc) based on equation VRc = 0.025 * TCI – 14.35 (Robison et al., 2000). Columns 3 and 4: Correlation between Acritarch Colour Alteration Index (AAI) of Legall et al. (1981) and Williams et al. (1998) and vitrinite reflectance (VRr) after Duggan and Clayton (2008, Fig. 4). Columns 5–7: Correlation between Spore Colour Index (SCI) of Fisher et al. (1980), Thermal Alteration Index (TAI) of Staplin (1969), Pollen/Spore Color “Standard” (SCS) of Pearson (1990) and vitrinite reflectance (VRr) after Marshall and Yule (1999, Fig. 31.1). Column 8: Shows approximate correlation between Index of Alteration of Microfossils Colour/ Colour Index of Microphytofossils (AMC/ICM) of Rovnina (1981, 1984) with miospore colour scales in Columns 5–7. Column 9: Correlation between graptolite reflectance (GRr) and vitrinite reflectance (VRc) based on equation VRc = 10(Log10GRr)/1.1 * 100.04/1.1(Bertrand, 1990a, Table 1). Column 10: Correlation between graptolite reflectance (GRr) and vitrinite reflectance (VRc) based on equation VRc = 0.9376 * GRr + 0.0278 (Bertrand and Malo, 2012). Column 11: Correlation between graptolite reflectance (GRr) and vitrinite reflectance (VRc) based on equation VRc = 0.73 * GRr + 0.16 (Petersen et al., 2013). Column 12: Correlation between chitinozoan reflectance (CRr) and vitrinite reflectance (VRc) based on equation VRc = 10(Log10CRr)/1.081 * 100.004/1.081 (Bertrand, 1990a, Table 1). Column 13: Correlation between chitinozoan reflectance (CRr) and vitrinite reflectance (VRc) based on equation VRc = (CRr – 0.08)/1.152 (Tricker et al., 1992). Column 14: Correlation between chitinozoan reflectance (CRr) and vitrinite reflectance (VRc) based on equation VRc = 0.8873 * CRr + 0.0124 (Bertrand and Malo, 2012). Column 15: Correlation between scolecodont reflectance (SRr) and vitrinite reflectance (VRc) based on equation VRc = 10(Log10SRr)/1.47 * 100.19/1.47 (Bertrand, 1990a, Table 1). Column 16: Correlation between (Bertrand and Malo, 2012). Column 17: Correlation between hydroid reflectance scolecodont reflectance (SRr) and vitrinite reflectance (VRc) based on equation VRc = 1.2038 * SR0.6824 r (HRr) and vitrinite reflectance (VRc) based on equation VRc = 0.6493 * HRr + 0.2126 (Bertrand and Malo, 2012). Coal stages and phases of oil and gas generation are mainly based on Taylor et al. (1998, Fig. 3.40 and 8.1). Correlation between Tmax and vitrinite reflectance (VRr) based on equation Tmax = 63.03 VRr + 389.11 (Petersen et al., 2013).


different types of vitrinites differs by about 0.1%, i.e. hydrogen enriched vitrinites containing resinous material within cell fillings show lower reflectance than telovitrinite without such hydrogen-rich material. Because these different types of vitrinites can be easily distinguished in coals and because the measurements are restricted to collotelinite, standard deviations are usually low and reproducibility is high, whereas higher standard deviations are observed in clastic sedimentary rocks and carbonates. Vitrinite reflectance analysis is a powerful tool to decipher maximum burial depth and temperature; in combination with fission track data and numerical petroleum system modelling, complex burial and uplift histories can be reconstructed. At high maturity levels the primary differences between vitrinite precursor materials are evened out as the material becomes more and more aromatic. On the other hand, organic particles become strongly anisotropic at the anthracite rank or at the anchimetamorphic stage of diagenesis. At this level, deformation of vitrinite modifies the temperature-induced changes of the chemical and optical properties. Especially, VRmax or bireflectance values of vitrinites in cleavage domains can be misleading being much higher than those in undeformed sandstones. On the other hand, rotational reflectance of vitrinites can be used as an important tool to decipher the stress and strain history of rocks at these high levels of diagenesis. For vitrinites as well as for most other macerals and microfossils, detailed geochemical analyses should be undertaken on well established maturity series in order to further understand chemical structure and chemical evolution. Advanced analytical micro-techniques (e.g. μ-FTIR, synchrotron radiation-based μ-FTIR, μ-Raman spectroscopy, X-ray absorption near edge structure spectromicroscopy, laser-pyrolysis–gas chromatography/mass spectrometry) might add to our understanding. Although the temperature-controlled colour change of conodonts is generally well understood other aspects have largely been neglected so far. Besides the work by Marshall et al. (2001) no further study was undertaken in which a systematic comparison of the chemical and structural evolution of contact metamorphosed conodonts is investigated (except those of increasing apatite crystal size which corresponds to increasing metamorphism). Generally, the CAI temperature estimations provided by Epstein et al. (1977) and Rejebian et al. (1987) based on the Arrhenius diagram need a critical review. Furthermore, there is a need to study systematically, whether an impact of strain has an influence on organic matter sealed within the conodont elements and if CAI 5 really defines the brittle–ductile boundary. In this context, further statistical studies on the average grain size of conodont elements (e.g. Burnett et al., 1988; Kovács and Árkai, 1987; Wiederer et al., 2002) and its correlation with other metamorphic parameters are necessary in order to establish more precise CAI temperature ranges in high temperature regimes. It seems necessary to compare optical properties of different types of organic particles especially vitrinite, solid bitumen, conodonts, graptolites and palynomorphs (Fig. 26). In addition, alteration of the macromolecular composition of graptolites and palynomorphs (except miospores) during thermal maturation is even nearly unexplored and hence a systematic comparison of the chemical and structural evolution is urgently needed. Finally, it is recommended to combine different maturity parameters in studies on thermal history and petroleum generation of sedimentary rocks, because information based on one parameter alone can be misleading. Within the oil window, optical maturity parameters can often be well combined with organic geochemical parameters, especially up to the peak oil generation stage. At higher maturity levels (high grade diagenesis, anchimetamorphism), when only few organic geochemical parameters can be used, optical parameters and clay mineral parameters such as illite and chlorite crystallization are particularly useful. Acknowledgements We are grateful to Alfred Traverse and Natalia Zavialova for providing us with the unpublished thesis of A. Hagerman Reed and the

111

Russian publications by L.V. Rovnina, respectively. Walter Riegel kindly commented on an earlier version of the manuscript. We would like to thank Philippe Janvier, Martin Langer and Jörg Maletz for their discussions and advice on the recent progress in ichthyolith, foraminifera and graptolite research. Slides with arthropod remains were kindly provided by Olaf Gosny. The help of Ulrike Lux and Jörg Schardinel in producing the figures is gratefully acknowledged. The authors thank Ivana Sýkorová and an anonymous reviewer for their comments and suggestions that greatly helped to improve the manuscript. Appendix A. Supplementary material Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.coal.2015.06.005. References Ainsworth, N.R., 1987. Pliensbachian Ostracoda from the Fastnet Basin, offshore Southwest Ireland. Geol. Surv. Irel. Bull. 4, 41–62. Ainsworth, N.R., Burnett, R.D., Kontrovitz, M., 1990. 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