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Paleointensity determination of Late Cretaceous basalts in northwest South Korea: implications for low and stable paleofield strength in the Late Cretaceous
© The Society of Geomagnetism and Earth, Planetary and Space Sciences (SGEPSS); The Seismological Society of Japan; The Volcanological Society of Japan; The Geodetic Society of Japan; The Japanese Society for Planetary Sciences; TERRAPUB. 2013
Received: 19 June 2013
Accepted: 27 September 2013
Published: 6 December 2013
To clarify geomagnetic field behavior in the Late Cretaceous, paleointensity and rock-magnetic studies were carried out on basalts with K-Ar ages of 78.6 ± 2.5 and 73.1 ± 1.6 Ma in northwest South Korea. A total of 314 samples (262 from six lava flows and 52 from massive basalts) were subjected to Thellier-type IZZI paleointensity experiments. Through the application of seven paleointensity selection criteria that mainly verify the thermal alteration of magnetic minerals and the stability of remanence direction, fifteen paleointensity data were obtained from five of the lava flows. Based on rock-magnetic experiments and microscopic observations, it is revealed that the measured paleointensity is carried by single-domain (titano)magnetite. The site-mean paleointensities ranged 13.1−21.3 μT, which corresponds to a virtual axial dipole moment of 2.3−3.8 × 1022 Am2. The combination of this result with selected data from the IAGA paleointensity database establishes the existence of different dipole moments according to rock type. In particular, the geomagnetic field strength recorded in Late Cretaceous crystalline volcanic rock was relatively low and stable, with a mean dipole moment of 4.0 ± 1.9 × 1022 Am2 regardless of geomagnetic field reversals.
The Earth’s magnetic field is generated by electrical currents flowing in the liquid outer core and is sustained permanently by a self-exciting geodynamo (e.g., Glatzmaier et al., 1999). The stability of the field is known to be controlled by the thermal structure in the core-mantle boundary, based on the analysis of the statistical properties of marine magnetic anomaly time scales (McFadden and Merrill, 1984) and a numerical simulation of the geodynamo (Olson et al., 2010). McFadden and Merrill (1984) proposed that long periods of geomagnetic polarity without any reversals are related to vigorous thermal activity at the base of the mantle. In contrast, Olson et al. (2010) suggested that periods of frequent change in geomagnetic polarity may be linked to heat flow increase at the core-mantle boundary. Thus, information about the variation in the strength of the Earth’s magnetic field over geologic time periods can provide valuable insights into the evolution of Earth as a planet. In the last few decades, numerous studies have tried to collect more detailed information on the variation in paleofield intensities. Paleointensity data, however, are still insufficient in terms of temporal and spatial resolution to be reliable (see the IAGA absolute paleointensity [PINT] database; Biggin et al., 2010) because of the scarceness of suitable materials and magnetomineralogical alterations during experiments.
One of the most dramatic changes in the geomagnetic field since the Mesozoic occurred in the Late Cretaceous. After a long period of normal geomagnetic polarity (the Cretaceous Normal Superchron, CNS, ~38 Myr), the reversal rate of the geomagnetic field began to increase from zero to 1.08 reversals per million years until ~60 Ma (McFadden and Merrill, 1997). The directional records of the geomagnetic field are well documented for that time, but the characteristics of paleointensity are not as well known, even though the paleointensity during the Late Cretaceous plays an important role in understanding the Earth’s evolution in relation to geomagnetic field reversal and heat flow across the core-mantle boundary. It is notable that the PINT database (Biggin et al., 2010) presents only four paleointensity studies (Juárez et al., 1998; Smirnov and Tarduno, 2003; Tauxe, 2006; Shcherbakova et al., 2007) satisfying modern standards (i.e., the use of the double-heating method with alteration checks) within the time interval 83–72 Ma. Of these, three studies used submarine basaltic glass (SBG) as an experimental material: Juárez et al. (1998) reported a high dipole moment of 9.0 ± 0.7 × 1022 Am2 at ~72 Ma; Tauxe (2006) suggested 11.3 ± 0.7 × 1022 Am2 at ~72 Ma and 6.8 ±5.1 × 1022 Am2 at ~77.6 Ma; and Smirnov and Tarduno (2003) documented a low dipole moment of 2.1 ± 0.7 × 1022 Am2 at ~76 Ma. The only study to have obtained paleointensity data from basalts (Shcherbakova et al., 2007) showed a low dipole moment of 3.3 ± 0.1 × 1022 Am2 at ~77.6 Ma. Given that the time interval of 83–72 Ma has been suggested to correspond with a stationary regime for the geodynamo and a low geomagnetic reversal rate (Gallet and Hulot, 1997), high dipole moments are plausible, as suggested by numerical simulation (e.g., Olson and Hagee, 1990). However, such a discrepancy between the dipole moments of various studies could be explained by the materials used for the experiments, given that the magnetic minerals in SBG are sometimes chemically altered during paleointensity experiments (Heller et al., 2002; Smirnov and Tarduno, 2003). Hence, paleointensity data obtained from magnetically stable volcanic rocks are needed to clarify geomagnetic field behavior in the Late Cretaceous.
In this study, paleointensity and rock-magnetic analyses were performed on Late Cretaceous basalts in northwest South Korea to provide more detail on paleointensity during the Late Cretaceous. In addition, paleointensity variation between different rock types is discussed using data from the PINT database.
The basalts are characterized by a dark grey color with a dominance of plagioclase and clinopyroxene, and are classified as tholeiitic basalt (Lee et al., 1983). For the pale-ointensity estimate, a total of 268 cores were collected from eleven sites: one site within each of the six layers of the lava flows (182 cores) and five sites within the massive basalts (86 cores). The sampling locations are shown in Fig. 1(b). The cores were oriented in-situ using an orientation device equipped with both a magnetic compass and a sun compass. In the laboratory, the cylindrical cores were cut into standard-sized samples (25 mm in diameter and 22 mm in length).
Paleointensity experiments were carried out for 314 samples using the in-field/zero-field and zero-field/in-field (IZZI) experimental protocol (Tauxe and Staudigel, 2004; Yu et al., 2004) of the modified Thellier-type double-heating method (Thellier and Thellier, 1959), with partial thermoremanent magnetization (pTRM) checks. The pTRM checks were performed at every second heating step to detect magnetomineralogical changes produced by heating throughout the paleointensity experiment. The details of the experimental procedure for pTRM checks have been described in several studies (e.g., Coe et al., 1978; Selkin and Tauxe, 2000; Tauxe and Staudigel, 2004). The IZZI protocol consists of two double-heating procedures (infield/zero-field [IZ] and zero-field/in-field [ZI]), and the IZ procedure alternates with the ZI procedure for every other temperature step. Compared to other methods, the IZZI method has an advantage in checking remanence stability (Yu and Tauxe, 2005) by testing an equivalence of blocking and unblocking temperatures without additional heating steps in a zero-field condition (e.g., McClelland and Briden, 1996; Riisager and Riisager, 2001). To impart TRM on the samples during each in-field step, a laboratory field (Blab = 25 μT in this study) was applied along a sample’s Z-axis. Thermal treatment was performed using the MMTD-80 thermal demagnetizer from Magnetic Measurements, and the remanence was measured using a Molspin spinner magnetometer. The temperature steps were 200°C, 300°C, 350°C, 400°C, 450°C, 500°C, 520°C, 540°C, 560°C, 570°C, 585°C, and 600°C. In some cases, additional temperature steps of 550°C and 555°C were applied because the remanence rapidly unblocked within the temperature interval 540–560°C. At every heating step, the set temperature was reproducible within <2°C. The experimental results are presented on an Arai plot (Nagata et al., 1963), in which the natural remanent magnetization (NRM) remaining after the zero-field step is plotted against the corresponding pTRM gained in the Blab for successive heating steps. In an ideal case, the ratio of NRM lost (Manc) to pTRM gained (Mlab) remains constant throughout the entire temperature interval. For each sample, the ratio of Manc/Mlab is obtained from the slope of the best-fit line for the NRM-pTRM data points within a selected temperature interval. The Earth’s ancient magnetic field (Banc) is then calculated as Banc = Manc/Mlab × Blab. The anisotropy of an anhysteretic remanent magnetization (AARM) tensor was also obtained for eight selected samples following the experimental procedure suggested by Selkin et al. (2000). To check the effects of remanence anisotropy on Banc, vector components for the measured remanences during the pa-leointensity experiments were corrected using the AARM tensor.
To identify magnetic carriers, thermal demagnetization of 3-axis isothermal remanent magnetization (IRM) (Lowrie, 1990) was conducted for 33 representative samples (three for each sampling site). Magnetizing fields of 2.5 T, 0.6 T, and 0.12 T were applied to the Z-, Y-, and X-axis respectively of a sample using an ASC Scientific IM-10–30 impulse magnetizer. The sample was then thermally demagnetized up to 700°C over 19 steps using an ASC Scientific TD-48 thermal demagnetizer. In addition, the temperature dependence of a magnetic susceptibility (k-T) curve was measured using a Bartington MS2 susceptibility-temperature system with a MS2WF furnace. The heating and cooling rate of the furnace was ~3.5°C/min. The Curie temperature (Tc) was determined from the point with the maximum gradient on the heating curve. In order to diagnose the domain state of magnetic minerals, hysteresis parameters at room temperature, including saturation remanent magnetization (Mrs), saturation magnetization (Ms), and coercive force (Hc), were measured using a Molspin vibrating sample magnetometer with a maximum magnetic field of 1 T. The coercivity of remanence (Hcr) was determined from the pattern of the hysteresis loop, where the reverse field reduces the remanence to zero, using PmagPy software (Jackson et al., 1990; Tauxe, 2010). After the magnetic measurements, electron microscopic observation was carried out for the selected samples using a Hitachi S-4800 FESEM integrated with energy dispersive X-ray spectroscopy to identify the morphological features and the chemical composition of the magnetic minerals in the samples.
4.1 Paleointensity selection criteria
Summary of the seven criteria for selecting paleointensity.
Number of points (Np)
To obtain reliable paleointensity statistics
Maximum angular deviation (MAD)
To ensure stability of remanence direction
Deviation angle (DANG)
To guarantee remanence direction toward the origin
Scatter parameter (σ/ǀbǀ)
To assure straightness of the selected data points
Fraction of the NRM ( f )
To avoid under- and over-estimated paleointensity
pTRM check (CKerror)
To ensure little influence of thermal alteration
Quality factor (q)
To ensure high-quality paleointensity data
4.2 Paleointensity determinations
Figure 4 shows examples of the Arai plot for rejected paleointensity results. Figure 4(a) displays an example of a sample that exhibits non-linear and zigzag patterns of data points on the Arai plot with a high σ/ǀbǀ (= 0.492) and CKerror (= 8.1%) at relatively low temperature steps (200–540°C). This result is probably caused by a difference in the capacity for acquiring TRMs between the IZ and ZI procedures (Yu et al., 2004; Yu and Tauxe, 2005). Figure 4(b) provides an example of unstable rema-nence direction (MAD = 22° and DANG = 18.4°) with a failed pTRM check (CKerror = 11.2%) at the high temperature interval of 450–585°C, despite a linear trend (σ/ǀbǀ = 0.054). On the vector diagram (inset in Fig. 4(b)) in particular, the remanence exhibits unstable demagnetizing behavior from 550°C to 585°C with a high angular deviation (MAD = 16.7°). Consequently, experimental results showing this behavior are rejected, though data points at the high temperature interval (450–585°C) are sufficiently linear to determine an acceptable paleointensity value. Figure 4(c) shows the importance of the pTRM checks; if there were no checks during the paleointensity experiments, there would be no other grounds for rejection. In the temperature steps above 500°C, the pTRMs increase gradually (i.e., CKerror(540°C) = 6.7% and CKerror(570°C) = 12.5%, Fig. 4(c)). Finally, Fig. 4(d) is an example of samples showing a short linear trend of data points on the Arai plot (f = 0.39) at a low temperature interval (200–500°C). Although a linear segment exists at the low temperature interval (grey dashed line), the remanence with a low unblocking temperature is likely to be contaminated by viscous remanent magnetization. In addition, the remanence direction is unstable above 500°C (inset of Fig. 4(d)) and accompanied by failed pTRM checks (CKerror = 9.5%).
Paleointensity results obtained from basalts in Yeoncheon.
Paleointensity selection criteriaa
VADMmσm (1022 Am2)
5. Rock-magnetic Properties
5.1 Magnetic mineralogy
5.2 Magnetic granulometry
6. Scanning Electron Microscope Observation
The lower panels in Fig. 7 show BSE images of iron oxides from the samples rejected in paleointensity determination. For the rejected samples from the lava flows, the iron oxides are characterized by secondary precipitation forming rims of iron oxide surrounding iron-rich silicate minerals (e.g., olivine) with overall grain sizes of ~20 μm (Fig. 7(d)). Such a large grain size may negatively affect the chance of obtaining reliable paleointensity estimates. In the case of the massive basalts, iron oxides predominantly occur in dendritic and cruciform with grain sizes of <10 μm (Fig. 7(e)). In particular, iron oxide adjacent to voids has a trellis texture (Fig. 7(f)). This texture is typical of hematite pseudomorphed after magnetite (i.e., mar-tite) when exposed to oxidizing conditions for long periods of time (e.g., Heider and Dunlop, 1987). The existence of hematite in samples from massive basalts has also been observed in earlier rock-magnetic investigations. Consequently, it is possible that the samples that were failed in paleointensity determination suffered from magnetominer-alogical alteration after cooling of the magmatic body or during thermal treatment in the paleointensity experiments.
Overall, observations using a scanning electron microscope reveal that the predominant magnetic phase in samples from the lava flows is Ti-poor titanomagnetite exhibiting a SD or PSD state, and which is expected to preserve primary remanent magnetization. In contrast, a considerable amount of the iron oxide in samples from the massive basalts suffers from secondary alteration. In particular, samples for which paleointensity was able to be determined are dominated by SD (titano)magnetite, observed in anhedral to subhedral form.
7. Paleointensity during the Late Cretaceous
It is known that the geometry of the Earth’s magnetic field has been predominantly dipolar for at least the past 250 Myr (e.g., Bloxham, 2000). Therefore, the measured intensity of the geomagnetic field should be expected to vary depending on geographic location. To compare the field intensities from various regions, an estimated pale-ofield should be converted into a virtual dipole moment (VDM) using magnetic paleolatitude derived from mean inclination (Smith, 1967). When directional information is uncertain, a VDM can be substituted with a virtual axial dipole moment (VADM) using the site latitude or estimated paleolatitudes from an apparent polar wander path (Barbetti et al., 1977). In the present study, we calculate a VADM instead of a VDM using the current latitude of the sampling sites, based on previous paleomagnetic results (Doh and Piper, 1994; Zhao et al., 1999; Doh et al., 2002; Kim et al., 2009) suggesting that the Korean Peninsula experienced clockwise rotation around its vertical axis with respect to Eurasia until the early Paleogene without significant latitudinal movement (i.e., a possible latitudinal difference of <0.5°).
In order to provide details for the paleointensity behavior during the Late Cretaceous (100–65 Ma), the six site-mean paleointensities in this study are compared with previously published paleointensity data archived in the 2012 version of the PINT database. Of the >250 entries from the database, 132 data sets from 18 different studies were selected for comparison. The selected data were all obtained using the Thellier-type method with pTRM checks, and were unrelated to a geomagnetic polarity transition. Figure 8 shows the selected Late Cretaceous VDMs or VADMs alongside our data. Four previously reported studies (Juárez et al., 1998; Smirnov and Tarduno, 2003; Tauxe, 2006; Shcherbakova et al., 2007) have investigated basalts of a similar age to those in this study (~80–70 Ma). Basalts from ~77.6 Ma in Georgia (Shcherbakova et al., 2007) had a VDM of 3.3 ±; 0.1 × 1022 Am2, consistent with the VADMs (= 2.3–3.8 × 1022 Am2) of this study. In contrast to the stable VDMs reported from basalts, results from SBG of ~77.6 Ma (Tauxe, 2006), ~76 Ma (Smirnov and Tarduno, 2003), and ~72 Ma (Juárez et al., 1998; Tauxe, 2006) reveal a mean VDM of 6.8 ±;5.1 × 1022 Am2, 2.1 ±; 0.7 × 1022 Am2, and 10.7 ±; 1.3 × 1022 Am2, respectively. Such a high fluctuation could be caused by the locking process of the geomagnetic field in SBG, which can cool too fast and thus not average out the secular variation. Some studies (e.g., Heller et al., 2002; Bowles et al., 2011) also argue that SBG is suspected to be altered during thermal treatments above 300°C. Indeed, Smirnov and Tarduno (2003) reported that their VDMs were under-estimated due to high temperature (above 350–400° C) mineral alteration. The VDMs of Smirnov and Tarduno (2003) are thus excluded from the paleointensity analysis of the Late Cretaceous.
Mean virtual (axial) dipole moment during the Late Cretaceous for each rock type.
Age interval (Ma)
VDMorVADM (× 1022 Am2)
Crystalline volcanic rock a
, , , this study
Submarine basaltic glass
Plutonic rocks (gabbro and dolerite)
Single plagioclase crystal
Sedimentary rocks of baked contact
All rock types
5.8±;3.5 (5.1±;1.8) –, , , this study
–, , , this study
Dipole moments from plutonic rocks in the Late Cretaceous are also documented in the PINT database. For the period of 100–83 Ma, the mean dipole moment is calculated to be 6.2 ±; 2.4 × 1022 Am2, which corresponds to a value between the mean dipole moments of crystalline volcanic rock and SBG. For the period of 83–65 Ma, only a single VDM (= 4.0 × 1022 Am2) is reported from Perrin et al. (1999), and accord with the mean dipole moment obtained from crystalline volcanic rock (Table 3). As the cooling rate of plutonic rock is much slower than during laboratory cooling, correction for the cooling rate is required (e.g., Dodson and McClelland-Brown, 1980; Fox and Aitken, 1980; Halgedahl et al., 1980; Walton and Williams, 1988). However, a commonly applied correction method has not yet been established.
Additional data from single plagioclase crystals (VDM of ~95 Ma = 12.7 ±; 0.7 × 1022 Am2; Tarduno et al., 2002) and from sedimentary rock of baked contact (VDM of ~86.4 Ma = 4.5 ±; 1.1 × 1022 Am2; Shcherbakova et al., 2008, 2009) are available within the time interval of 100–83 Ma (Table 3) in the PINT database. The mean VDM from single plagioclase crystals is almost three times higher than that from crystalline volcanic rock (= 4.0 ±; 2.1 × 1022 Am2), while the mean VDM from sedimentary rock of baked contact is comparable to that from crystalline volcanic rock. However, the absence of paleointensity data from single plagioclase crystals and sedimentary rock in the later part of the Late Cretaceous (83–65 Ma) makes it difficult to discuss overall geomagnetic field variation during the Late Cretaceous.
The mean dipole moments of entire rock types (= 6.7 ±; 3.9 × 1022 Am2 for 100–83 Ma; = 5.8 ±; 3.5 × 1022 Am2 for 83–65 Ma) are also higher than that of crystalline volcanic rock, which is caused by the higher values from SBG, plutonic rock, and single plagioclase crystals. Nevertheless, the dipole moments are lower than the present field strength (Table 3). It is obvious that there exists a degree of variability when comparing different materials (e.g., compare crystalline rock, SBG, or single plagioclase crystals during the 100–83 Ma period; Table 3). On the other hand, geomagnetic field intensity during the Late Cretaceous was relatively low and stable, when each material is considered separately or all materials are considered collectively. In order to check for consistency from a more rigorously selected data set, a second group of paleointensity data was selected from the PINT database by applying two additional criteria (e.g., Biggin et al., 2009): (1) at least three samples should be used to calculate the VDM or VADM; and (2) the standard deviation should not be greater than 25% of the mean paleointensity. The calculated mean dipole moments from the second group of paleointensity data for each time interval and rock type are given within parentheses in Table 3. When the two sets of mean dipole moments are compared, it is noticeable that little difference exists between them, and that the characteristic Late Cretaceous feature of a relatively low, stable geomagnetic field can also be found in the second data set (Table 3). Considering that the reversal rate started to increase after the CNS (Fig. 8), it can be speculated that the initiation of geomagnetic field reversal was triggered without significant changes in average pa-leofield strength (see Fig. 8 and Table 3). This interpretation is supported by Goguitchaichvili et al. (2004), who suggested that there is no relationship between reversal rate and the strength of the geomagnetic field.
In this study, absolute paleointensity values for Korean basalts dating from the Late Cretaceous—after the Cretaceous Normal Superchron (CNS)—are obtained. The estimated site-mean paleointensities range from 13.1 to 21.3 μT, which corresponds to virtual axial dipole moments (VADMs) of 2.3−3.8 × 1022 Am2. Based on the radiometric age of the basalts, the VADM for ~73.1 Ma is calculated to be 3.1 ±;0.7 × 1022 Am2. According to the results of rock-magnetic investigations and scanning electron microscope observations, it is revealed that Ti-poor titanomagnetite (either SD or PSD grain size) is the major magnetic carrier in the lava flows, while hematite or hematite with magnetite is dominant in the massive basalts. In particular, samples for which paleointensity was able to be determined were predominantly composed of SD (titano)magnetite, observed in anhedral to subhedral form. A compilation of Late Cretaceous geomagnetic dipole moments obtained from crystalline volcanic rock and submarine basaltic glass had in common no significant time-dependent variation up until at least 65 Ma, regardless of geomagnetic field reversals. Based on this geomagnetic field behavior during the Late Cretaceous, it can be speculated that the initiation of geomagnetic field reversal after the CNS occurred without significant changes in average paleofield strength.
This research was supported by the Basic Science Research Program through the National Research Foundation of Korea (NRF), funded by the Ministry of Science, ICT and Future Planning (MSIP, 2011-0009679). This manuscript was greatly improved by constructive comments and suggestions by Dr. G. A. Paterson and an anonymous reviewer, as well as by the journal Editor Prof. T. Yamazaki.
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